Geology and Geochemistry of Epithermal Systems - Reviews Vol 2-Search

REVIEWS IN ECONOMIC GEOLOGY

Volume 2

GEOLOGY AND GEOCHEMISTRY OF

EPITHERMAL SYSTEMS B. R. Berger a P. M. Bethke, Editors

CONTENTS

THE GEOTHERMAL FRAMEWORK FOR EPITHERMAL DEPOSITS R . W. Henley A PRACTICAL GUIDE TO THE THERMODYNAMICS OF GEOTHERMAL FLUIDS

AND HYDROTHERMAL ORE DEPOSITS R . W. Henley & K . L . Brown THE BEHAVIOR OF SILICA IN HYDROTHERMAL SOLUTIONS R . 0. Fournier CARBONATE TRANSPORT AND DEPOSITION IN THE EPITHERMAL ENVIRONMENT R . 0. Fournier FLUID INCLUSION SYSTEMATICS IN EPITHERMAL SYSTEMS R . J . Bodnar, T. J . Reynolds, & C . A . Kuehn LIGHT STABLE-ISOTOPE SYSTEMATICS IN THE EPITHERMAL

ENVIRONMENT C. W. Field & R . H . Fifarek GEOLOGIC, MINERALOGIC, AND GEOCHEMICAL CHARACTERISTICS OF

VOLCANIC-HOSTED EPITHERMAL PRECIOUS-METAL

DEPOSITS D . 0. Hayba, P. M . Bethke, P. Heald, & N . K . Foley GEOLOGIC CHARACTERISTICS OF SEDIMENT-HOSTED, DISSEMINATED

PRECIOUS-METAL DEPOSITS IN THE WESTERN UNITED STATES W. C . Bagby & B . R . Berger RELATIONSHIP OF TRACE-ELEMENT PATTERNS TO ALTERATION AND

MORPHOLOGY IN EPITHERMAL PRECIOUS-METAL DEPOSITS M . L . Silberman & B . R . Berger RELATIONSHIPS OF TRACE-ELEMENT PATTERNS TO GEOLOGY IN HOT-SPRING

TYPE PRECIOUS-METAL DEPOSITS B . R . Berger & M . L. Silberman BOILING, COOLING, AND OXIDATION IN EPITHERMAL SYSTEMS: A NUMERICAL

MODELING APPROACH M . H . Reed & N. Spycher USING GEOLOGICAL INFORMATION TO DEVELOP EXPLORATION STRATEGIES

FOR EPITHERMAL DEPOSITS S . S . Adams

Series Editor: James M. Robertson

SOCIETY OF ECONOMIC GEOLOGISTS

REVIEWS IN ECONOMIC GEOLOGY (ISSN 0741-0123)

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REVIEWS IN ECONOMIC GEOLOGY

(ISSN 074141123) Volume 2

GEOLOGY AND GEOCHEMISTRY OF

EPITHERMAL SYSTEMS

ISBN 0-9613074-1-2

Volume Editors:

B. R. BERGER P. M. BETHKE Branch of Exploraton Geochemistry Branch of Resource Analysis

U. S . Geological Survey U. S . Geological Survey MS 973 MS 959, National Center

Box 25046, Federal Center Reston, VA 22092 Denver, CO 80225-0046

Series Editor: JAMES M . ROBERTSON New Mexico Bureau of Mines & Mineral Resources Campus Station Socorro, NM 87801


The Authors:

Samuel S. Adams 3030 Third Street Boulder, CO 80302

William C. Bagby Branch of Western Mineral Resources U. S . Geological Survey MS 901 345 Middlefield Road Menlo Park, CA 94025

B. R. Berger Branch of Exploration Geochemistry U.S. Geological Survey MS 973 Box 25046, Federal Center Denver, CO 80225-0046

Philip M. Bethke Branch of Resource Analysis U.S. Geological Survey MS 959, National Center Reston, VA 22092

R. J . Bodnar Department of Geological Sciences Virginia Polytechnic Institute and State University Blacksburg, VA 2046 1

K. L. Brown Chemistry Division D.S.I.R., Private Bag Taupo New Zealand

Cyrus W. Field Department of Geology Oregon State University Corvallis, OR 9733 1-5506

Richard H. Fifarek Department of Geology Southern Illinois University Carbondale, IL 62901

N. K. Foley Branch of Resource Analysis U .S . Geological Survey MS 959, National Center Reston, VA 22092

Robert 0 . Fournier Branch of Igneous and Geothermal Processes U.S Geological Survey MS 910 345 Middlefield Road Menlo Park, CA 94025

Daniel 0 . Hayba Branch of Resource Analysis U.S. Geological Survey MS 959, National Center Reston, VA 22092

Pamela Heald Branch of Resource Analysis U.S. Geological Survey MS 959, National Center Reston. VA 22092

R. W. Henley Chemistry Divsion D.S.I.R.. Private Bag Taupo New Zealand

C. A. Kuehn Department of Geosciences The Pennsylvania State University University Park, PA 16802

Mark H. Reed Department of Geology University of Oregon Eugene, OR 97403

T. J. Reynolds FLUID, Inc. P.O. Box 6873 Denver, CO 80206

M. L. Silberman Branch of Exploration Geochemistry U. S . Geological Survey MS 912 Box 25046, Federal Center Denver, CO 80225-0046

N. Spycher Department of Geology University of Oregon Eugene, OR 97403

GEOLOGY & GEOCHEMISTRY OF EPITHERMAL SYSTEMS

CONTENTS

F O R E W O R D . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . x

PREFACE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xi BIOGRAPHIES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . x v i

CHAPTER 1

THE GEOTHERMAL FRAMEWORK OF EP1THERMA.L DEPOSITS R . W . Henley

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION. 1

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . HYDROTHERMAL SYSTEMS IN GENERAL 1

Collision-Related Amagmatic Hydrothermal Systems Terrestrial Magma-Related Hydrothermal Systems

. . . . . . . . . . . . . . . . . . . . . . . . TERRESTRIAL MAGMATIC-HYDROTHERMAL SYSTEMS 4

Laree Scale Structure -. - - -

Natural Discharges Hydrothermal Eruption Vents Heat and Mass Flow in Geothermal Systems

. . . . . . . . . . . . . . . . . . . . . . . . . . . CHEMISTRYOFGEOTHERMALDISCHARGES. 11

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . EPITHERMAL ORE-FORMING SYSTEMS 12

Requirememts for Ore Deposition Chemistry of Systems Responsible for Ore Formation Chemical and Physical Processes in Ore Formation Host-Rock Relations

SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . EPILOGUE. 21

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ACKNOWLEDGMENTS 21

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . REFERENCES 21

CHAPTER 2

A PRACTICAL GUIDE TO THE THERMODYNAMICS OF GEOTHERMAL FLUIDS AND HYDR0THERMA.L ORE DEPOSITS

R . W . Henley and K . L . Brown

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION.. 25

. . . . . . . . . . . . . . GEOLOGICAL CHARACTERISTICS OF THE BROADLANDS GEOTHERMAL SYSTEM 25

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . FLUID CHEMISTRY 26

FLUID-MINERAL EQUILIBRIA: ALTERATION MINERALOGY . . . . . . . . . . . . . . . . . . . . . 28 . . . . . . . . . . . . . . . . . . . . . FLUID-MINERAL EQUILIBRIA: TRACE-METAL CONTENTS 32

Lead . Gold . Other Metals: Copper, Silver, and Arsenic

MINERAL DEPOSITION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 Silica Calcite Metal Sulfides and Gold

ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 REVIEW QUESTIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 APPENDIX . . . . . . . . . . . a s . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43

CHAPTER 3

TEE BEHAVIOR OF SILICA IN EYDROTEERMAL SOLUTIONS R . 0 . Fournier

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION 45

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . SOLUBILITIES OF SILICA MINERALS 45

. . . . . . . . . . . . . . . . . . THE BEHAVIOR OF DISSOLVED SILICA IN HOT-SPRING SYSTEMS 46

ALKALINE WATERS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 ACIDWATERS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 REACTIONWITHGLASS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 AMORPHOUS SILICA-CHALCEDONY RELATIONS . . . . . . . . . . . . . . . . . . . . . . . . . . 51

. . . . . . . . . . . . . . . . . . . . . . SPECULATIONS REGARDING SOME TEXTURES OF QUARTZ 51

Jasperoid and Massive Replacement of Limestone by Silica Quartz Solubility at High Temperatures

CONCLUSIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 APPENDIX . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60

CARBONATE TRANSPORT AND DEPOSITION IN TEE EPITBERMAL ENVIRONMENT R. 0 . Foumier

INTRODUCTION. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63

C02 DISSOLVED IN AQUEOUS SOLUTIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63

THE SOLUBILITY OF CALCITE IN AQUEOUS SOLUTIONS . . . . . . . . . . . . . . . . . . . . . . 67

SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71

FLUID-INCLUSION SYSTEMATICS IN EPITEERMAL SYSTEMS R. J. Bodnar, T . J . Re,ynoZds, and C . A. Kuehn

INTRODUCTION. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 73

INFORMATION AVAILABLE FROM FLUID-INCLUSION PETROGRAPHY . . . . . . . . . . . . . . . . . . 73

IDENTIFICATION OF FLUID INCLUSIONS TRAPPED FROM BOILING SOLUTIONS . . . . . . . . . . . . 79

IDENTIFICATION OF GASES IN FLUID INCLUSIONS FROM THE EPITHERMAL ENVIRONMENT . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83

INTERPRETATION OF FLUID INCLUSIONS FROM THE EPITHERMAL ENVIRONMENT . . . . . . . . . . . . 93

APPLICATION OF FLUID INCLUSIONS IN EXPLORATION FOR EPITHERMAL PRECIOUS-METAL DEPOSITS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 94

SUGGESTIONS FOR FUTURE FLUID-INCLUSION RESEARCH . . . . . . . . . . . . . . . . . . . . . 95

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96

CHAPTER 6

LIGar STABLE-ISOTOPE SYSTEMATICS IN THE EPITHERMAL ENVIRONMENT C . W . Field and R. H . Fifarek

INTRODUCTION. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 CONVENTIONS, SYSTEMATICS, AND RATIONALE . . . . . . . . . . . . . . . . . . . . . . . . . 99

Fractionation Equilibrium Reaction Applications

GEOLOGIC DISTRIBUTIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 110

Hydrogen and Oxygen Carbon Sulfur

EPITHERMALDEPOSITS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113

Hydrogen and Oxygen

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . SUMMARY 124

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . REFERENCES 125

GEOLOGIC, MINERALOGIC, AND GEOCHEMICAL CHARACTERISTICS OF VOLCANIGEOSTED EPITBERMAL PRECIOUS-METAL DEPOSITS

D. 0. Hayba, P. M . Bethke, P . Heald, and 1. K. Foley

SUMMARY OF THE CHARACTERISTICS OF VOLCANIC-HOSTED EPITHERMAL ORE DEPOSITS . . . . . . . . 129 Characteristics of Adularia-Sericite-Type Deposits Characteristics of Acid-Sulfate-Type Deposits Summary of Characteristics

THE ADULARIA-SERICITE ENVIRONMENT: CREEDE AS AN EXAMPLE . . . . . . . . . . . . . . . . . 136 Creede as an Exemplar Summary of Important Studies Geologic and Mineralogic Characteristics Geochemical Environment Hydrologic Environment Boiling and Mixing in the Ore Zone Summary of Creede Mineralization

. . . . . . . . . . . . . . . . . THE ACID-SULFATE ENVIRONMENT: SUMMITVILLE AS AN EXAMPLE 151

Geologic and Yineralogic Characteristics Geochemical Environment - - - - ~ ~ -

Summary of Summitville Mineralization

GEOTHERMAL INTERPRETATION OF VOLCANIC-HOSTED EPITHERMAL DEPOSITS . . . . . . . . . . . . . 158 Adularia-Sericite Deposits Acid-Sulfate Deposits

MECHANISMS OF ACID-SULFATE ALTERATION . . . . . . . . . . . . . . . . . . . . . . . . . . 159 ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 162

CHAPTER 8

GEOLOGIC CHARACTERISTICS OF SEDIMENT-HOSTED, DISSEMINATED PRECIOUS-METAL DEPOSITS IN THE WESTERN UNITED STATES

W. C. Bagby and B. R. Berger

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INTRODUCTION. 169

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CLASSIFICATION 169

REGIONAL GEOLOGIC CHARACTERISTICS OF DEPOSITS IN MINERAL TRENDS AND ISOLATED DEPOSITS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172

The Getchell Trend The Carlin Trend

ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 228 REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 230

CHAPTER 1 0

RELATIONSHIPS OF TRACE-ELEMENT PATTERNS TO GEOLOGY IN HOT-SPRINGTYPE PRECIOUS-METAL DEPOSITS

B . R. Berger and M . L . SiZberman

CONTROLS ON TRACE-ELEMENT PATTERNS . . . . . . . . . . . . . . . . . . . . . . . . . . . . 233 TRACE-ELEMENT PATTERNS IN STUDIED DEPOSITS . . . . . . . . . . . . . . . . . . . . . . . . 235

Hasbrouck Mountain, Nevada Round Mountain, Nevada

DISCUSSION. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 245

CHAPTER 11

BOILING, COOLING, AND OXIDATION IN EPITAERMAL SYSTEMS: A NIlMERICAL MODELING APPROACH M . H. Reed and N . F . Spycher

BOILING . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249 BOILING RESULTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252 DISCUSSION OF BOILING AND COOLING . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252

Sulfide and Carbonate Mineral Precipitation Precipitation of Silicates Boiling Without Fractionation and Cooling Only

SUPER- AND SUB-ISOENTHALPIC BOILING . . . . . . . . . . . . . . . . . . . . . . . . . . . 258 BOILING AND GOLD PRECIPITATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 261 THE HOT-SPRING ENVIRONMENT . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 262

Condensation of the Boiled Gas Oxidation of Gases to Produce Acid-Sulfate Waters Reaction of Gases with Meteoric Ground Water Gold Precipitation from Mixing of Acid-Sulfate Water with Boiled Aqueous Phase

Gold Precipitation from Mixing of Oxygenated Ground Water with Boiled Aaueous Phase

SUMMARY . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 269 ACKNOWLEDGMENTS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 270

CHAPTER 12

USING GEOLOGICAL INFORMATION TO DEVELOP EXPLORATION STRATEGIES FOR EPITEIERMAL DEPOSITS

S . S. Mums

INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 273

SOME CONSIDERATIONS IN THE USE OF GEOLOGICAL INFORMATION IN EXPLORATION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 273

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . STRATEGIC FACTORS 274

Organizational Objectives Commodity Prices Financial Resources Exploration Organization Regulations and Land Availability Competitor Activity Previous Ex~loration

Risk - . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . HUMAN FACTORS 279

Personal Objectives Education and Training Problem Solving Intuition and Creativity Uncertainty Aversion to Loss

DEVELOPMENT OF MINERAL-DEPOSIT MODELS . . . . . . . . . . . . . . . . . . . . . . . . . . 282

Organization of Geologic Information Model Terminology Level of Model Development

DATA-PROCESS-CRITERIA MODEL . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 286

Definition of a Mineral-Deposit Type Compilation of Analog Deposits Selection of Geologic Data Data-Process Linking Identification of Formation Processes Evaluation of Data-Process Links Selection of Diagnostic Criteria Evaluation of Data-Process-Criteria Model Application of Data-Process-Criteria Model Summary of Data-Process-Criteria Model

Exploration

CONCLUSIONS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 296

REFERENCES . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 297

TABLE OF CONVERSION FACTORS . . . . . . . . . . . . . . . . . . . . . . . . . . . Inside Back Cover

FOREWORD

Geology and Geochemistry of Epi thermal Systems--Volume 2 of Reviews Economic Geology--was c r e a t e d t o accompany a Socie ty of Economic Geologists (SEG) short course of t h e s a m e n a m e t h a t was given in October , 1985, prior t o t h e annual meetings of t he Geological Society of America and Associated Societies in Orlando, Florida. As was t h e case with Volume 1, t he final published version of Volume 2 unfortunately postdates t h e short course by some months.

Geology and Geochemistry of Epi thermal Systems presents a synthesis of t h e current understanding of t h e processes responsible for t he concentra t ion of me ta l s (especially gold and silver) in near-surface environments, provides an overview of t h e sys temat ics of t h e most important approaches t o t h e study of epi thermal ores and processes, and summarizes t h e geology of both sediment-hosted and volcanic-hosted epi thermal precious-metal deposits.

Af t e r t h e volume editors, t h e most significant contributors t o t h e production of this voIume were t h e members of t h e Editorial Support Group, Branch of

Exploration Geochemistry, U.S. Geological Survey, Denver, Colorado. These ladies, Marilyn A. Billone, Candace A. Vassalluzzo, and especially Pamela S. D e t r a and Dorothy B. Wesson, accomplished t h e long, arduous, and of ten f rus t ra t ing job of assembling, edit ing, and format t ing t h e book with a uniformly high level of professionalism and good cheer. Their e f fo r t s a r e gratefully acknowledged. Carol Hjellming of t h e New Mexico Bureau of Mines and Mineral Resources (NMBMMR) editing s taf f checked, balanced, and helped in terpre t t h e chemical equations; Lynne McNeil (NMBMMR) fo rma t t ed t h e cutlines. Lastly, I wish t o express my continuing appreciation t o t h e New Mexico Bureau of Mines and Mineral Resources and i t s Director, Frank Kottlowski, who provide t h e Series Editor with t ime, space, and encouragement.

J a m e s M. Robertson Ser ies Editor Socorro, NM March, 1986

PREFACE

In a speech on May 10, 1911, before t h e Geological Society of Washington, Waldemar Lindgren described his sys temat ic classification of a l l types of mineral deposits. One of his ca tegor ies included deposits re la ted t o intrusive and eruptive igneous rocks t h a t form veins a t shallow depths t h a t contain open- cavity filling t ex tu res and t h a t have been a primary source of "bonanza" grades of gold and silver--the epi thermal deposits. Historically, most of t h e ores in epi thermal sys tems have been mined f rom quar tz veins, breccias, or disseminations t h a t a r e associated with non-marine volcanic rocks. Open-space filling textures and s t ructures a r e common--comb s t ructure , crustification, symmetr ica l banding, and crystal-lined vugs. Ore minerals include nat ive gold, nat ive silver, e lec t rum, argent i te , sulfosalts, tellurides, and selenides and of ten t h e common sulfides sphalerite, galena, and chalcopyrite. Common gangue minerals a r e quar tz , adularia, calcite, bar i te , rhodochrosite, and fluorite. Alteration is commonly widespread in epithermal systems, particularly in t h e upper portions of the vein systems; among the a l tera t ion phases a r e quartz, adularia, illite, chlorite, alunite, and kaolinite.

Lindgren (1928) recognized the difficulty of developing a rigid subsidiary classification scheme for epi thermal deposits; he separa ted them into six categories:

Confidence

I. Cold deposits 2. Argentite-gold deposits 3. Argent i te deposits 4. Gold selenide deposits 5. Gold telluride deposits 6 . Gold telluride deposits with a luni te

Nolan (1933) and Ferguson (1929) f e l t t h a t few of these six character is t ics were res t r ic ted enough to be diagnostic and proposed only two classes of epithermal sys tems based on t h e weight ra t io of gold t o silver, silver-gold, and gold-silver. Based on his experience with deposits in Nevada, Ferguson (1929) found t h a t t he re is a bimodal distribution of gold-silver ratios, and Nolan (1933) f e l t t h a t t h e bimodality was due t o genet ic processes.

For t h e silver-gold deposits, Nolan (1933) noticed t h a t through-going faul t f issures control t h e o re and f e l t t h a t this implies a deep origin for t he source of t h e metals. Nolan (1933) also noted t h a t t he precious- me ta l ores a r e very commonly sharply l imited above and below by approximately parallel surfaces referred t o a s t h e o re "horizon." He suggested tha t these limits a r e re la ted t o temperature . Base meta ls tend t o increase a t and below t h e base of t he lower surface of t h e precious-metal ore. Figure 1 i s a longitudinal, ver t ica l projection of the Last Chance-Confidence silver-gold vein in t h e Mogollon mining district , New

n Last Chance

700-FT LEVEL

500 1000 FEET 900-FT LEVEL

Figure P.1. Vertical, longitudinal projection of the Confidence-Last Chance vein in the Mogollon mining district, New Mexico (Ferguson, 1927). Banded quartz vein is continuous along strike with ore grade material occurring in specific masses (stippled areas) in the vein. The tops and bttoms of the silver-rich ore bodies describe near parallel surfaces referred to as the "ore horizon."

Mexico (Ferguson, 1927) i l lus t ra t ing the o re horizons, t h e shape of o r e bodies, and t h e typical distribution of o r e grades within a continuous banded quartz-adularia- ser ic i te vein. Burbank (1933) reported t h a t base me ta l s appear t o be m o r e abundant in silver-gold deposits in regions of sedimentary rocks with overlying volcanic rocks and in thick, volcanic sequences with a long history of volcanic activity. In contras t t o the silver-gold deposits, Nolan (1933) noted t h a t gold- silver deposits a r e commonly within or close t o small, shallow intrusive bodies a n d t h a t t h e ore-controlling f r ac tu re sys t ems a r e f requent ly more discontinuous than those associa ted with silver-gold deposits. The gold-silver o res a r e also m o r e irregular in distribution than the silver-gold ores. Nolan f e l t t h a t this irregularity may be r e l a t ed t o the complex thermal regimes in these types of sys tems due t o t h e shallow intrusive activity. Figure 2a shows a ser ies of plan views of t h e January mine, Goldfield mining dis t rc t ,

Nevada and a cross section through the January sha f t (Ransome, 1909) showing t h e relationships of o re t o quartz-alunite-kaolinite replaced wallrock ("ledge mat ter") and t h e host rocks. Figure 2b shows t w o cross sect ions f rom Ransome (1909,. p. 154) of t h e Combination mine in Goldfield ~ l lus t r a t ing t h e irregular ver t ica l distribution of bonanza-grade o re masses within t h e "ledge matter." Also, t h e ore bodies were not pers is tent along strike.

Although Waldemar Lindgren (1928) recognized the corre la t ion between epi thermal systems and a c t i v e geothermal sys tems, i t was Donald E. White (1955, 1981) who championed the deta i led study of ac t ive sys tems and t h e application of t h e results and concepts derived f rom these studies t o epi thermal ore deposits. The impac t of White's leadership in t h e study of hydrothermal systems, in general, and epi thermal systems, in particular, was recognized by t h e Society of Economic Geologists when i t held a symposium in

JANUARY SHAFT

109-FT LEVEL 51-FT LEVEL 8 I -FT LEVEL

80-Ft Leve l

130-Ft Leve l

180-Ft Leve l

230-Ft Leve l

280-Ft Leve l

160-FT LEVEL 232-FT LEVEL 2 8 3 - F T LEVEL

380-Ft L e v e l

JANUARY SHAFT a

232-FT LEVEL

283-FT LEVEL

CROSS SECTION

0 5 0 I q O 150 FEET I I I

Figure P.2. a). Plan views of the January mine at selected mining levels and a cross section through the January shaft Goldfield mining district, Nevada (Ransome, 1909). Bonanza-grade gold ores occur in replaced dacite referred to as "ledge matter." The ore bodies are not persistent either down- dip or along strike, and occur both on the hanging wall side of the ledge and on the foot wall side. b). Two cross sections from Ransome (1909, p. 154) of the Combina- tion mine in the Goldfield district. So- nanza-grade gold ores occur in irregular, discontinuous masses within the ledge. The ledge follows a lithologic contact and flattens with depth.

xii

his honor in February, 1984 entit led: Geothermal Systerns and Ore Deposits. I t clearly emphasized the value of using act ive geothermal a reas a s models of fossil, ore-forming hydrothermal systems.

Thus, the evolution of understanding of t h e geology and genesis of epi thermal precious-metal deposits has followed a pathway from t h e early, vividly descriptive studies of mining districts such a s the Comstock Lode, Nevada (Becker, 18821, Cripple Creek, Colorado (Lindgren and Ransome, 19061, and Waihi, New Zealand (Bell and Fraser , 1912) t o the la ter , topical studies on s t ruc tu re (Wisser, 1960), a l tera t ion (Hernley and Jones, 19641, s table isotopes (Taylor, 19731, and fluid chemis t ry (Barton e t al., 1977). The most recent research on epi thermal deposits has built on these past studies and has emphasized the thermal and compositional roles of volcanic rock terranes; t h e genesis, significance, and pat tern of a l tera t ion mineralogies; t h e sources of t h e geothermal fluids and t h e paleohydrology of the systems; and, t he chemical conditions surrounding t h e deposition of t h e o r e minerals.

The present volume is an a t t e m p t t o provide a synthesis of t he cu r ren t s t a t e of geological and geochemical knowledge of epithermal precious-metal systems. I t follows on, and should be used in conjunction with, t he f i rs t volume in this series: Mineral-Fluid Equilibria in Hydrothermal Systems by Henley e t al. (1984). In the present volume we have a t t empted to provide a framework for understanding the sys temat ics of controls on fluid compositions and of me ta l and gangue t ranspor t and deposition. The s t ructure , dynamics, and transport properties of ac t ive geothermal sys tems a r e used a s a s tar t ing point. With ac t ive sys tems a s a reference , t he evolution of fluid compositions and t h e const ra in ts on me ta l and gangue transport and deposition in t h e epitherrnal environment a r e explored. The sys temat ics of fluid inclusion and light stable-isotope applications is developed because these two approaches have been so useful in the development of our understanding of epi thermal processes. The importance of boiling, cooling, and oxidation in t ranspor t properties of epi thermal sys tems is evaluated through a numerical modelling approach. With the foregoing a s background, t h e observational base and i t s in terpre ta t ion for epi thermal ore deposits in continental volcanic and sedimentary ter ranes is explored through summaries of the geologic, mineralogical, and geochemical character is t i s of, and trace- e lement distributions in, some well-studied epi thermal o re deposits. The final chapter is devoted t o t h e use of our understanding of epithermal sys tems in the development of exploration strategies.

This volume does not a t t e m p t t o be the final word on epi thermal o r e deposits, nor does i t claim comprehensive t r ea tmen t . The absence of a chapter on t h e hydrology of epi thermal systems documents t h e f a c t t h a t our cu r ren t understanding of this aspect is woefully inadequate. I t does not r e f l ec t a lack of recognition of the importance of hydrologic controls. Similarly, this volume focuses on volcanic- and sediment-hosted epi thermal deposits in the cordillera of western North America, particularly t h e United States. I t does not t r e a t aspects of alkaline- or basaltic-rock re la ted deposits such a s Cripple Creek,

Colorado, and Vatacoula, Fiji, nor does i t t r e a t t h e relationship of epi thermal sys tems t o deeper hydrothermal sys tems responsible for t he format ion of porphyry-type deposits. Again our reason is t h e lack of an adequa te observational base. Our primary purpose in organizing this volume and t h e r e l a t ed Short Course has been t o s t imulate cr i t ica l studies t o improve our cu r ren t understanding of epi thermal deposits and processes ra ther than t o document i t . Perhaps our omissions will serve this purpose equally a s well a s ou r inclusions.

REFERENCES

Barton, P. B., Jr., Bethke, P. M., Roedder, E., 1977, Environment of ore deposition in the Creede mining district , San Juan Mountains, Colorado: 111. Progress toward in terpre ta t ion of the chemis t ry of t h e ore-forming fluid fo r t h e OH vein: Economic Geology, v. 72, p. 1-25.

Becker, G. F., 1882, Geology of t h e Comstock lode and the Washoe district: U.S. Geological Survey Monograph 3, 442 p.

Bell, J. M., and Fraser , C., 1912, The g rea t Waihi gold mine: New Zealand Geological Survey, Bulletin - 15.

Burbank, W. S., 1933, Epithermal base-metal deposits in Ore deposits of t h e Western States: American Ins t i tu te of Mining Metallurgical Engineers, New York, P a r t VI, p. 641-652.

Ferguson, H. G., 1927, Geology and o re deposits of t he Mogollon mining district , New Mexico: U.S. Geological Survey Bulletin 787, 100 p.

Ferguson, H. G., 1929, The mining dis t r ic ts of Nevada: Economic Geology, v. 24, p. 131-141.

Hemley, J . J., and Jones, W. R., 1964, Chemical aspects of hydrothermal a l tera t ion with emphasis on hydrogen metasomatism: Economic Geology, v. 59, p. 538-569.

Henley, R. W., Truesdell, A. H., and Barton, P. B., Jr., 1984, Fluid-mineral equilibria in hydrothermal systems: Society of Economic Geologists, Review in Economic Geology, v. 1, p. 267.

Lindgren, W., 1928, Mineral Deposits: Third Edition, McGraw Hill, New York, 1049 p.

Lindgren, W., and Ransome, F. L., 1906, Geology and gold deposits of t h e Cripple Creek district , Colorado: U.S. Geological Survey, Professional Paper 54, 516 p.

Nolan, T. B., 1933, Epithermal precious-metal deposits in Ore deposits of t h e Western States: American Ins t i tu te of Mining Metallurgical Engineers, New York, P a r t VI, p. 623-640.

Ransome, F. L., 1909, The geology and o r e deposits of Goldfield, Nevada: U.S. Geological Survey Professional Paper 66, 258 p

Taylor, H. P., Jr., 1973, 18/616/0 evidence for meteoric-hydrothermal a l tera t ion and ore deposition in t h e Tonopah, Comstock Lode, and Goldfield mining districts, Nevada: Economic Geology, v. 68, p. 747-764.

White, D. E., 1955, Thermal springs and epi thermal o re Wisser, E., 1960, Relation of o r e deposition t o doming deposits: Economic Geology, F i f t i e th Anniversary in the North American Cordillera:. Geological Volume, p. 99-154. Society of America, Memoir 77.

White, D. E., 1981, Ac t ive geothermal sys tems and hydrothermal o r e deposits: Economic Geology, Seventy-fifth Anniversary Volume, p. 392-423.

ACKNOWLEDGMENTS

As is t rue for any e f fo r t of t h e scope of th is volume, many people in addition t o t h e edi tors played key roles along t h e road t o final publication. The t i m e and e f fo r t expended by each author is great ly apprecia ted a s a r e t h e contributions of t h e large cadre of individual reviewers who have offered insights and a l ternat ive perspectives t o t h e authors. Technical support t o t h e edi tors including manuscript preparation and revision, final format t ing for publication, and badgering of both editors and authors was provided by t h e Editorial Support Group, Branch of Exploration Geochemistry, U.S. Geological Survey. Within this group we would especially like t o thank Pamela Detra ,

Dorothy Wesson, Marilyn Billone, and Candy Vassalluzzo. An ear l ier version of th is t e x t was assembled for use a t t h e Socie ty of Economic Geologists Short Course by t h e Branch of Exploration Geochemistry Cler ica l Support Group. Finally, we would like t o express appreciation fo r t h e patience of J amie Robertson, Ser ies Editor, Reviews in Economic Geology, and t h e suppor t of t h e Society of Economic Geologists.

Byron R. Berger Philip M. Bethke

BIOGRAPHIES

BYRON R. BERGER received a B.A. degree in ~ c o n o m i c ~ e o l o ~ ~ f rom Occidental College in 1966 and a M.S. in Geology f rom the University of California, Los Angeles in 1975. He worked a s a petroleum exploration geologist for Standard Oil Company of California f rom 1968-1970 and a minerals exploration geologist and research scientist for Cont inenta l Oil Company from 1971-1977. He joined the U.S. Geological Survey in 1977, and has been involved in research on epi thermal precious-metal deposits and t h e relationship of magma genesis t o o re genesis. He is current ly t h e Chief of t h e Branch of Exploration Geochemistry. He is an adjunct assistant professor of geology in t h e Depar tment of Geological Sciences a t t h e University of Colorado, Boulder, where he has taught courses on t h e geology and geochemistry of epi thermal ore deposits and exploration geochemistry. H e is a member of several professional societies including the Geological Society of America and t h e American Geophysical Union.

PHILIP M. BETHKE received a B.A. degree in Geology -- from Amherst College in 1952 and a Ph.D. in Geology (specialization in Mineralogy and Ore Deposits) f rom Columbia University in 1957. He was Assistant Professor of Geology a t t h e Missouri School of Mines and Metallurgy (now t h e University of Missouri-Rolla) from 1955 to 1959. He joined the U.S. Geological Survey a s a WAE research geologist in 1957 and t ransferred t o full t i m e in 1959. His research has combined field and laboratory approaches t o t h e study of hydrothermal ore deposits, particularly t o epithermal vein systems. He has held several administrative positions with the U.S.G.S., most recently, Chief of t h e Branch of Experimental Geochemistry and Mineralogy. He is a member of several professional societies and is current ly a Councillor of t he Society of Economic Geologists. He has been act ive in t h e establishment of t h e SEG Short Course Series, and is currently Chairman of the Short Course Commit tee .

SAMUEL S. ADAMS received B.A. and M.A. degrees from Dartmouth College in 1959 and 1961, and a Ph.D. degree from Harvard University in 1967. From 1964 to 1977 he served a s mine neolonist, exploration - - . geologist, exploration manager, and exploration vice president, employed by International Minerals and ~ h e m i c a i ~ o ; ~ o i a t i o n - and then the Anaconda Company. During this period, his work emphasized sediment-hosted mineral deposits, particularly potash and uranium. Since 1977 he has served a s a lec turer and consultant t o industry, research organizations, and

government agencies in t h e a reas of mineral deposits, exploration, and resource assessment. His principal research in t e re s t is t h e representation of d a t a and concepts fo r a l l types of mineral deposits in coherent and predictive models for exploration and resource studies. He is currently a Councillor of t he Society of Economic Geologists and the Geological Society of America.

WILLIAM C. BABGY received a Ph.D. degree in Ear th Science f rom t h e University of California, Santa Cruz, in 1979 based on pet rogenet ic research of Ter t iary volcanic rocks in t h e Sierra Madre Occidental , Mexico. His industry experience includes geologic evaluation of volcanic-hosted uranium in the McDermit t ca ldera complex, Nevada, and t h e bulk mineability potential of t he amythest silver vein sys tem a t Creede, Colorado. Industry research included development of an occurrence model fo r hot spring-related gold deposition based on the McLaughlin gold deposit in California. Present research in teres ts a r e focused on t h e genet ic aspects of sediment-hosted precious-metal deposits.

ROBERT J, BODNAR received an M.S. degree from t h e University of Arizona and a Ph.D. degree f rom The Pennsylvania S t a t e University and has been involved in various a spec t s of fluid-inclusion research for t h e past 10 years. He worked for 1 year a s a research geochemist in t h e Ore Deposits Group of Chevron Oil Field Research Company and is currently an ass is tant professor in t h e Depar tment of Geological Sciences a t Virginia Polytechnic Insti tute and S t a t e University.

KEVIN BROWN received an M.S. degree in Chemist ry -- in 1969 and a Ph.D. degree in Chemical Crystallography in 1972 from t h e University of Auckland, New Zealand. Except for a two-year sojourn a t t h e E.T.H. in Zurich, he has worked a t t h e Depar tment of Scientific and Industrial Research, New Zealand. Initially in Wellington, his research in t e re s t cen te red around the crys ta l s t ructures of organic reaction in termedia tes , but he gradually c a m e down t o e a r t h with t h e crys ta l s t ructures of some new epi thermal minerals. In 1981, he shifted t o the Geothermal Section a t Wairakei, where his present research is concerned with exper imenta l studies of mineral deposition from geothermal fluids.

CYRUS W. FIELD received a B.A. degree in Geology --- f rom Dartrnouth College in 1956 and M.S. and Ph.D. degrees in Economic Geology, Geochemistry, and Petrology f rom Yale University in 1957 and 1961,

respectively. H e worked a s a n exploration geologist during the summers of 1955, 1956, and 1957 for t h e Oliver Iron Mining Company and Quebec Car t i e r Mining Company subsidiaries of the U.S. Steel Corporation, and served a s a research geologist from 1960 t o 1963 with t h e Bear Creek Mining Company division of Kennecot t Copper Corporation. In 1963, he joined the facul ty of Oregon S t a t e University where he is currently Professor of Geology. His research in teres ts a r e largely concerned with the geology and geochemistry of hydrothermal mineral deposits; particularly t h e application of s table isotope and major-minor-trace e l emen t investigations t o thei r genesis. He is a member of several professional societies and was Vice President of t h e Society of Economic Geologists in 1981.

RICHARD H, FIFAREK received a B.S. degree in Geology from t h e University of Washington in 1974, and M.S. and Ph.D. degrees in Geology (specialization in Economic Geology) f rom Oregon S t a t e University in 1982 and 1985, respectively. From 1974 t o 1984, he worked periodically a s an exploration geologist (4 yrs.) for severa l mining companies, a s a research assistant/ scientist (1 yr.) a t t h e facil i t ies of t he Branch of Isotope Geclogy (Denver), U.S. Geological Survey, and a s an ins t ructor fo r Oregon S t a t e University. Presently, he is a n ass is tant professor in t h e Depar tment of Geology a t Southern Illinois University where he teaches and conducts research in economic geology and isotope geochemistry. His research in teres ts include in tegra ted geologic (field) and geochemical investigation of massive sulfide and epi thermal Au-Ag deposits, and modeling t h e isotopic evolution of fluids and rocks in hydrothermal systems.

NORA K. FOLEY received a B.S. degree in Geology --- and Mineralogy f rom t h e University of Michigan in 1978 and an M.S. degree in Geological Sciences f rom Virginia Polytechnic Ins t i tu te and S t a t e University in 1980. She is currently working towards a doctoral degree in Geology through Virginia Polytechnic Insti tute and S t a t e University. Since 1980, she has been a research geologist a t t he U.S. Geological Survey in Reston, Virginia. Her research has included fluid-inclusion and isotopic studies of d i f ferent types of ore deposits, including Ag- and base-metal-bearing, epi thermal deposits, sediment-hosted, stratabound, Pb-Zn deposits, and Kuroko-type massive sulfides.

ROBERT O, FOURNIER received an A.B. degree in Geology in 1954 from Harvard College and a Ph.D. in Geology (specializing in Economic Geology, in general, and the Ely porphyry copper deposit, in particular) from the University of California a t Berkeley in 1958. Since then, he has been a research geologist with t h e U.S. Geological Survey. His research in teres ts have ranged from laboratory studies of mineral-water in teract ions a t hydrothermal conditions appropriate for shallow levels in the crus t , t o field studies of presently ac t ive hydrothermal systems, including Yellowstone National Park, Coso and Long Valley, California, and Zunil, Guatemala. Experimental studies have emphasized solubilities of

si l ica species in wa te r and saline solutions. H e has also been a leader in t h e development of several chemical geothermometers and mixing models t h a t a r e now widely used in t h e exploration for geothermal resources. His present research focuses mainly on internally consistent chemical, isotopic, and hydrologic models of presently ac t ive hydrothermal systems. He has served on NATO commi t t ees t o review geothermal energy development programs in Iceland, France, Greece , Portugal, and Turkey, and other commi t t ees t o review geothermal exploration programs in Argentina and Thailand. He was Chairman of the Organizing C o m m i t t e e fo r t h e 1975 United Nations International Symposium on Geothermal Energy, and Chairman of t h e Technical Program Commi t t ee for t h e 1985 GRC International Symposium on Geothermal Energy. He now serves on panels t o oversee geothermal developments in Costa Rica and Panama, and several U.S. Continental Scientific Drilling Commit tees . He is a member of several societies and has served on t h e Board of Directors of t he Geochemical Society and the Geothermal Resources Council.

DANIEL 0. HAYBA received a B.A. degree in Geology --- f rom the College of Wooster in 1976 and an M.S. degree in Geochemistry and Mineralogy f rom the Pennsylvania S t a t e University in 1979 following a study of t h e Salton Sea geothermal systern. From 1978 t o 1980, he worked for Exxon Production Research Company on computer modeling of ore deposits. Since t h a t t ime, he has been a research geologist with t h e U.S. Geological Survey where his research has been di rected towards understanding the ore-forming processes in epi thermal systems.

PAMELA HEALD received a B.A. degree in Geology in 1971 f rom Vassar College and an M.S. degree in Geology f rom George Washington University in 1977. She has been a research geologist a t t h e U.S. Geological Survey since 1972. Her research has included spect ra l ref lec tance and s t ructura l studies in Nevada, with a focus on ore deposits, and mineralogical and geochemical studies t o evaluate ore- forming processes in epi thermal precious- and base- me ta l deposits.

RICHARD W. HENLEY received a B.S. degree in Geology in 1968 from t h e University of London and a Ph.D. degree in Geochemistry from The University of Manchester in 1971 following experimental studies of gold t ranspor t in hydrothermal solutions and the genesis of some Precambrian gold deposits. He was Lecturer in Economic Geology Memorial University of Otago, New Zealand, from 1971 to 1975, and a t Memorial University, Newfoundland, until 1977. Research in teres ts have focused on the mode of origin of a number of different types of ore deposits including post-metamorphic gold-tungsten veins, porphyry copper, massive sulfide, and placer gold deposits. He is currently with the Geothermal Chemist ry Section of the Depar tment of Scientific and Industrial Research a t Wairakei, New Zealand, and a visiting lec turer a t t he Auckland Geothermal

Ins t i tu te . Through 1983-84, he was a Fulbright Fellow and Gues t Inves t iga tor a t t h e U.S. Geological Survey a n d during t h a t t i m e produced Volume I of th is Review series. His present research includes a number of i so tope and chemica l studies re la t ing t o t h e explora t ion and development of geothermal sys t ems and geo the rma l implications fo r t h e origin of o r e

C. A. KUEHN received a n M.S. degree f rom t h e --- Pennsylvania S t a t e University and has 7 years of exper ience in explora t ion fo r sediment-hosted gold deposits . H e i s cu r r en t ly a n NSF Reseaqch Ass is tant and Ph.D. cand ida t e a t t h e Pennsylvania S t a t e University and par t - t ime employee of t h e U.S. Geological Survey working on t h e Car l in gold deposit.

MARK H. REED received a B.A. deg ree in Chemis t ry --- and in Geology f rom Car l e ton College in 1971 and M.A. and Ph.D. deg rees in Geology a t t h e University of Cal i fornia , Berkeley, in 1977. His Ph.D. research was on t h e geology and geochemistry of t h e massive sulfide deposits of t h e West Shas ta District , California. F rom 1977 through 1979, he worked fo r t h e Anaconda Minerals Company a t But te , Montana. Since t h a t t ime, h e has t augh t and conducted research a t t h e University of Oregon, where he is current ly Associa te Professor of Geology. His research has focused on a l t e r a t ion and me ta l zoning in t h e porphyry copper and l a rge vein deposits a t B u t t e and t h e geochemis t ry of hydrothermal a l te ra t ion , me ta l t ranspor t , and o r e deposit ion in massive sulfide and epi thermal systems.

geologist specializing in t h e application of fluid inclusions t o mineral explora t ion fo r t h e pas t 5 years.

MILES L. SILBERMAN received a B.S. deg ree f rom t h e -- Ci tv Universitv of New York and M.S. and Ph.D. d e k e e s f rom t i e University of Roches ter , New York. H e is a member of t h e Branch of Exploration Geo- chemis t ry of t h e U.S. Geological Survey, with current assignments t o t h e Redding, Cal i fornia (CUSMAP) project , and t o t h e study of t h e geochemis t ry of volcanic and metamorphic-hosted gold deposi t s in t h e western U.S. and nor thern Mexico. Previous work fo r t h e U.S;G.S. included geochronological, geochemical, and regional geological s tudies of precious- and base- me ta l deposits in t h e G r e a t Basin and Alaska, and t ec ton ic syntheses with par t icular focus on t h e relationships of hydrothermal precious-metal deposits t o magmat i c and me tamorph ic evolution. Between tours a t t h e U.S.G.S., he designed and supervised exploration programs f o r precious-metal deposits in t h e G r e a t Basin fo r t h e Anaconda Minerals Company.

NICOLAS G S P Y C H E R received a B.S. deg ree in Ea r th Sciences in 1979 and a Dipl. e s Sc. in Exploration Geophysics in 1980 f rom t h e University of Geneva, Switzerland. He is now a Ph.D. candidate and research ass is tant a t t h e University of Oregon. His present research includes s tudies of t h e t r anspor t of arsenic and ant imony in hydrothermal solutions, t h e mixing proper t ies of geothermal gases, and t h e geochemical modeling of hot spring systems.

T. J. REYNOLDS received a n M.S. degree f rom t h e -- University of Arizona and has been a n explora t ion

xviii

Chapter 1 THE GEOTHERMAL FRAMEWORK OF EPITHERMAL DEPOSITS

R. W. Henley

INTRODUCTION

In the con tex t of exploration for epi thermal deposits, why study geothermal sys tems a t all? Af t e r all, not one exploited system t o d a t e has been shown by drilling t o harbor any economically significant m e t a l resource--but then until recent ly not one had been drilled for o ther than geothermal energy exploration.* The l a t t e r involves drilling t o depths of 500-3000 me te r s in search of high t empera tu res and zones of high permeabili ty which may sustain fluid flow t o production wells for s team separa t ion and e lec t r ic i ty generation. In many cases such exploration wells have discovered disseminated base-metal sulfides with some silver and argillic-propylitic a l tera t ion equivalent t o t h a t commonly associated with ore- bearing epi thermal systems (Browne, 1978; Henley and Ellis, 1983; Hayba e t al., 1985, this volume). In general, however, geothermal drilling ignores the upper f ew hundred me te r s of t h e ac t ive sys tems and drill s i tes a r e s i tuated well away from natura l f ea tu res such a s hot springs o r geysers, t h e very f ea tu res whose character is t ics (silica sinter, hydrothermal breccias) a r e recognizable in a number of epi thermal precious- m e t a l deposits (see, for example, White, 1955; Henley and Ellis, 1983; White, 1981; Berger and Eimon, 1983; Hedenquist and Henley, 1985a; and ear l ier workers such a s Lindgren, 1933). Knowledge of t h e upper f e w hundred me te r s of ac t ive geothermal sys tems is s can t and largely based on in terpre ta t ion of hot-spring chemistry. Tantalizingly, in a number of hot springs, t rans i tory red-orange precipitates occur which a r e found t o be o re grade in gold and silver and which ca r ry a sui te of e lements (As, Sb, Hg, TI) now recogpized a s character is t ic of epi thermal gold deposits (Weissberg, 1969).

*Kennecot t has recently announced significant gold discoveries in still ac t ive geothermal fields on Lihir and Simberi Islands, Papua, New Guinea.

Today's ac t ive geothermal sys tems occupy t h e s a m e tectono-volcanic niche a s those hydrothermal systems, preserved from the past, which hosted t h e near-surface (0-1000 m) formation of epi thermal o r e deposits in the Ter t iary volcanic t e r r anes of t he Circum-Pacific region and elsewhere--the relatively shallow origin of these deposits result ing in the i r loss by erosion from erstwhile similar, but older, terranes. Formed a t deeper levels (2-5 km or so) beneath calc- alkaline volcanoes in these same volcanic t e r r anes (Sillitoe, 1973; Henley and McNabb, 1978), porphyry-

type copper and molybdenum deposits a r e preserved in both Ter t iary and much older hydrothermal systems. The purpose of this chapter is t o review some of the principal chemical and physical character is t ics of t he ac t ive geothermal sys tems which a r e essential t o the understanding of t h e origin of epi thermal o r e deposits and therefore t o thei r successful exploration. For more detailed information, t h e reader i s referred t o the publications c i ted in t h e text .

HYDROTHERMAL SYSTEMS IN GENERAL

The t e rm "hydrothermal" encompasses a l l types of hot-water phenomena in t h e ear th ' s c rus t although most commonly the t e r m is used in r e fe rence t o those associated with impressive geyser activity, aes thet ica l ly a t t r ac t ive hot pools, e tc . These fea tures a r e most common in volcanic a r e a s such a s Yellowstone National Park, U.S.A., Iceland, or in the Taupo Volcanic Zone of New Zealand, but other t e r ranes also host hydrothermal ac t iv i ty even though subsurf a c e temperatures may be relatively low and surface f ea tu res less impressive. Warm springs in the Rocky Mountains, t h e European or New Zealand Alps, or in the sedimentary mass i fs of cen t r a l Europe a r e examples, and i t is clearly important for mineral exploration t o discriminate these types of systems f rom those in more favorable geological environments.

Geothermal sys tems a r e extraordinarily abundant in t h e tectonically ac t ive zones of t h e earth 's crus t and may be broadly classified according t o thei r p la te t ec ton ic se t t ing and principal source of heat (Table 1.1). Chemical differences a r i se from the sources of recharge water and contribution of gases from magmat ic or metamorphic sources. Warm springs also occur in t h e tectonically s table crus t where the deep crus ta l penetra t ion of groundwater occurs in favorable sedimentary format ions such a s l imestones and the hea t supply is t he ambient continental hea t flow.

Each of these classes of geothermal systems appears t o have some corre la t ive preserved in the geologic pas t and most commonly recognized a s one or another of t h e various families of hydrothermal ore deposits. For magma-related hydrothermal systems, these range from ophiolite-hosted massive sulfides through t h e polymetall ic massive sulfides of island a rc s t o t h e porphyry copper and epi thermal precious- me ta l deposits of ter res t r ia l cont inenta l terranes, while for amagmat i c sys tems these range f rom t h e Mississippi Valley and r e l a t ed base-metal deposits in sedimentary basins t o t h e post-metamorphic vein deposits associated with orogeny.

Tab le 1.1--Crustal s e t t i n g of hydrothermal sys t ems c l a s s i f i e d accord ing t o p r i n c i p a l heat -source and c r u s t a l h o s t .

CRUSTAL HOST/ HEAT SOURCE MAGMATIC AMAGMATIC

Ocean ic Ridge, h o t s p o t , back-arc b a s i n ----- Magmatic a r c

C o n t i n e n t a l C r u s t a l e x t e n s i o n (Hot s p o t , r i f t )

P l a t e c o l l i s i o n

P l a t e - i n t e r i o r b a s i n s

Collision-Related Amagmat ic Hydrothermal Systems

Only recent ly have d a t a become available f rom geothermal investigations in mountain belts. In the Southern Alps of New Zealand, for example, hot springs occur in t h e centra l , relatively aseismic region with t h e highest uplift r a t e (10-20 mm/year) where t h e combination of uplift and erosion "exposes" a thermal ant ic l ine with near-surface gradients up t o 150°C/km (Allis e t al., 1979). A similar environment is proposed for hot springs in o ther collision-related mountain belts. Recent drilling a t Yangbajing (Tibet) and in t h e Pa rba t i Valley (N. India), for example, has located hot waters up t o 1 7 0 ' ~ (Giggenbach e t al., 1983) which a r e predominantly me teo r i c in origin, but contain low 3 ~ e t o 4 ~ e ratios typical of helium of deep crus ta l origin.

The uplift s e t t i ng of these hydrothermal sys tems is perhaps analogous t o t h a t of L a t e Mesozoic post- metamorphic gold and scheel i te veins on t h e South Island (New Zealand) and, by inference, similar deposits in much older terranes. Examples a r e t h e gold veins of the Valdez Group (S. Alaska), Mother Lode (California), Yellowknife (Northwest Territories) and Kalgoorlie (W. Australia). In each of these, in contras t t o t h e epi thermal precious-metal depos' ts discussed below, vein quar tz is enriched in "0 rela t ive to host rocks. This f ea tu re has led many workers (e.g., Henley e t al., 1976; Fyfe and Kerrich, 1984) t o suggest a metamorphic origin for t h e hydrothermal fluid; vein format ion occurring f rom fluids of metamorphic dehydration origin in response t o post-metamorphic uplift and/or overthrusting. I t may also be possible, however, t o generate these same isotope character is t ics by in teract ion of me teo r i c wa te r and rocks a t a low wa te r t o rock ra t io opening the possibility t h a t such deposits may be much shallower* in origin than t h e 10 t o 20 km generally considered.

*In this paper the t e rm "shallow" is used ra ther irreverantly t o r e fe r t o depths less than about 500 meters. In o re deposit research, depths (es t imated perhaps f rom fluid-inclusion data) a r e generally also used irreverantly, taking no account of t he importance of topographic relief; of t h e order of *I00 m in silicic

volcanic ter ranes , i 1000 m in andesit ic volcanic t e r r anes and for t h e mountain bel t systems, 2000 m.

Terres t r ia l Magma-Related Hydrothermal Systems

By contras t , sys tems in volca ' c ter ranes have f'i high 3 ~ e t o 4 ~ e ra t ios and the 6 0 of a l tera t ion minerals are , with few exceptions, depleted re la t ive t o primary minerals. Tern e ra tu res encountered during g drilling range up t o 400 C (Batini e t al., 1983a) and waters a r e predominantly me teo r i c in origin, and typified by the presence of chloride ion with mC1- > >mSy= --they a r e here, for convenience, designated chlorl e waters. Although some highly saline fluids a r e evolved in r i f t zones such a s t h e Imperial Valley (California), salinit ies a r e typically low, clustering around 10,000 rng/kg C1 (1.6 wt.-% NaCl equivalent) in andesi t ic volcanic ter ranes , 1000 mg/kg in rhyolit ic volcanic t e r r anes and much lower in basalt ic volcanic terranes. Dissolved gas, always preponderantly C 0 2 , a f f e c t s a major contras t between sys tems and ranges from very low (0.01 wt.-% C 0 2 ) a t Wairakei (New Zealand) and Ahuachapan (El Salvador) t o several wt.-% a t Broadlands and Ngawha (New Zealand) (see Table 1.3). Other dissolved components a r e controlled by mineral-fluid and gas-gas reactions. Alteration assemblages in these types of geothermal sys tems correspond closely t o those encountered in epi thermal and porphyry-style mineral deposits.

The deep hydrologic s t ruc tu re of t h e ter res t r ia l geothermal sys tems is controlled by the convective upflow of chloride wa te r s (evolved by water-rock i magma interact ion a t depths of 5 to 8 km) but above depths of around I km surface topography plays a major ro le in t h e dispersion of the chloride water by introducing a l a t e ra l flow component toward topographic lows. Boiling occurs a s chloride water rises through t h e sys tem, t h e resultant s t eam migrating to the surface independently where near- surface condensation and oxidation of co-transported H2S produces sulfate-dominated s team-heated waters. These f ea tu res a r e incorporated in t h e general model of t he s t ruc tu re of a geothermal system reproduced in Figure l . l a (from Henley and Ellis, 1983).

adv a r g i l l i c a l t e r a t i o n w i t h Near

a Meteoric Acid neutrat pH water Sulphate & Chlor~de Dilute

a l t e r a t i o n Groundwater

- . . ., . 0 1 km

HEAT AND MASS (NaCI, C02, 502, Hz0 . ) TRANSFER FROM MAGMA SYSTEMS.

KEY

Pre-Volcanic Basement Steam-heated Acid ~ 0 ~ / ~ 0 ~ - r i c h waters

Intrusive Volcanics SO;- C l - waters ( f i g I b )

Low Permeability Stratum 0 Near neutral Chloride waters e.g. Mudstones (k i th in 200" Isotherm approx.)

Two Phase Region Water Liquid + Steam ( +Gas)

- ~-

Figure l.la. Generalized structure of a typical geothermal system in silicic-volcanic terrane. Notice the overall size of the system relative to the size of the discharge features (i.e., hot springs, etc.). The temperature distribution shown is based on the Wairakei system where a west-to-east flow occurs in the upper portion of the system and boiling occurs above about 500 meters. In other systems such as those in Figure 1.2, more or less lateral flow may occur. Boiling may extend to much greater depths if C02 contents are high (see text), and higher temperatures may occur at shallower depths than shown in this figure, as at Mokai (Fig. 1.2d).

The re la t ively high relief of andesi te volcanic t e r r anes resul ts in l a t e ra l f lows of hot chloride water for up t o 20 km while t h e occurrence of near-surface magmas exsolving gases (HCI, SO2, etc.) of ten produces high t empera tu re fumaroles and/or ac id sulfate-chloride c ra t e r lakes, such a s those on Mount Ruapehu, New Zealand (P la t e 1.1) and El Chichon, Mexico (Giggenbach, 1974; Kyosu and Kurahashi, 1984; Casadeval l e t al., 1984). These l a t t e r fea tures , with thei r associa ted in tense advanced argil l ic a l tera t ion, a r e possible corre la t ives of t he upper portions of t he type of hydrothermal sys tems responsible for t h e format ion of gold--(enargite) sulfide deposits of t he "Goldfield type" (Ransome, 1909) such a s Goldfield (Nevada), Summitvil le (Colorado), Bor (Yugoslavia), and elsewhere. Figure 1.lb provides a general s t ruc tu ra l model for th is geothermal environment. They m a y also be re la ted in some cases t o t h e upper portions of developing porphyry copper deposits (Sillitoe, 1983).

The geochemistry and s t ructure of magma- r e l a t ed hydrothermal sys tems have been reviewed in a number of r ecen t t e x t s t o which the reader i s r e fe r r ed for background reading and discussion of hydrothermal chemistry--see, for example, Ellis and Mahon, 1977;

Henley and Ellis, 1983; Henley e t al., 1984. A brief summary of hydrothermal chemis t ry i s given in Henley and Brown (1985, this volume). In t h e remainder of this chapter a t tent ion is focused on those a spec t s of t h e chemistry and s t ructure of geothermal sys tems re levant t o t h e understanding of t h e format ion of epi thermal ore deposits.

TERRESTRIAL MAGMATIC-HYDROTHERMAL SYSTEMS

Large-Scale St ructure

Early in the commercia l development of the Wairakei geothermal field in New Zealand, t he accumulating d a t a from exploration wells showed (a) t h a t t h e fluids present were= directly exsolved from shallow bodies of crystall izing magma and (b) t h a t t he hydrothermal ac t iv i ty seen a t t h e su r face was a minor phenomenon associated with the discharge of a very large, deeply convecting body of heated groundwater (Elder, 1966). Using analog and numerical modelling, Elder and other research sc ient is ts showed t h a t convection, with a depth sca l e of a t l e a s t 5 km, was

Meteoric water

Neutral chloride water

F i g u r e l.lb. S t r u c t u r e o f a t y p i c a l g e o t h e r m a l s y s t e m i n a n d e s i t i c - v o l c a n i c t e r r a n e s emphas iz ing ( 1 ) e x t e n s i v e lateral f l o w and (2) g e n e r a t i o n o f a large a d v a n c e d - a r g i l l i c a l t e r a t i o n zone i n r e sponse to h igh- l eve l volcanism.

(Modif ied and reproduced w i t h p e r m i s s i o n f rom Henley and E l l i s , 1983.)

Plate 1.1. Oblique, aerial view of the Waiotapu system, New Zealand from the southeast. Topo- graphic features may be related to the system map (Fig. L.2b). Mount Tarawera (on the horizon) is a composite rhyolite dome which, in 1886, violently erupted &salt through an axial rift. Associated phenomena were 6?e destruction of the Pi& and WLite Silica Terraces (Henley et al., 1984, Plate 1.21, and a n~mber of hydrothermal eruptiors in the btomahana-Waimanq geothermal system.

The natural discharge dominating the surface expression of the system is *Ae Pool (middle right) which occupies a hydrothermal eruption vent formed 900 years ago and which may overlie some 0.1 million ounces of gold formed 'by boiling in the conduit of the ,ml. Surface antimony-arsenic precipitates xcur which are ore-grade in silver and gold (photo D. L. Homer, N. Z. Geological Survey).

responsible for t h e e x t r e m e thermal gradients and t empera tu re pa t t e rns observed in the exploration drilling. A t t h e s a m e tinie, at Wairakei and in other fields, t he e f f e c t s of near-surface (depths less than 1000 m) s t ra t igraphy and s t ruc tu re and of relief- control led groundwater fiow became evident Iargely through geophysical techniques, especially resistivity surveying (Healy and Hochstein, 1973). Hanaoka (1983) has numerically modelled t h e e f f e c t s of topographic r e i i e l on near-surface hot-water flow and i t s dispersion by cold groundwater. This e f f e c t is partly responsible for t h e mushrooming of isotherms shown in many convective models and field cross sections.

Figure 1.2 provides some examples of t he iaterai-flow character is t ics and distribution of natura l discharges in a number of geothermal fields explored by drilling, with perhaps t h e Niokai field in New Zealand (Fig. 1 . 2 ~ ) being a particuiarly good il lustration of l a t e ra l flow as shown in cross-section in Figure 1.2d. In geothermal sys tems hosted by sil icic volcanic rocks, su r face topography is primarily control led by block-faulting o r ca ldera collapse providing relief of a f e w hundred me te r s and consequent l a t e ra l flow over d is tances of up t o about 5 km. In t h e higher relief t e r r a n e typical of andesi t ic volcanism, more e x t r e m e i a t e ra i flow occurs up t o about 20 km, An additional f e a t u r e of ac t ive andesi t ic volcanic t e r r anes is t he occurrence of high-level volcanism which allows volcanic gas t o vent t o summit fumaroles or t o summit c ra t e r lakes (Giggenbach, 1974) and t o maintain high-ievei "perched" aquifers containing very acid sulf ate-dominated waters. Exploration wells a t high eievation in such ter ranes o f t en encounter vapor-dominated geotherma! environments.

In the majority of systems, liquid water provides t h e continuum for fluid fiow but in other, f a r less common systems, water vapor dominates the discharges of deep exploration wells. The pre- exploitation s t a t e s of these "vapor-dominated" sys tems a r e poorly known and various models have been produced based on production d a t a from exploited fields. For example, fo r t h e Geysers (California) and fo r Eardereilo (Italy), White e t ai. (1971) suggest th.e presence of a very deep convecting brine overlain by an "alteration-sealed" cap of vapor. Of particular i n t e re s t is t h e association of these sys tems with epi thermai mercury and gold mineralization (e.g., McLaughlin, California), but both t h e Geysers and larder el!^ also contain base-metal sulfides and other "ore-related1' minerai phases in drill co re (Beikin et al., 1983; Sternfe!d, 1981) which suggest t h a t t h e present system has evo!ved from some previous liquik- dominated s t a t e . Others have suggested t h a t e ievated gas-content (dominantly C 0 2 ) perhaps coupled with relatively low host-rock porosity, may account fo r t he vapor-dominated cha rac te r o i well-discharges and post-exploitation pressure data. I t is in teres t ing t o n o t e t h a t most of t h e explored "liquid-dominant" geothermal systems, in si l icic volcanic t e r r anes especially, a r e associated with t ec ton ic subsidence (about -5 mm per year in the Taupo Volcanic Zone in New Zealand), but both the Geysers and Larderello occur in regions of high t ec ton ic uplift associated with volcanism. Quant i ta t ive d a t a f rom the Geysers region

a r e not available, although regional topography and erosion a r e suggestive of high uplift rates. A t Lardere l lo uplift r a t e s a r e of the order +5 t o 13 mm per year (M. Puxeddu, persona: communication^ and a r e evidenced by t h e coas t i ine migration of the Pisa area. The high hea t i low and geothermal ac t iv i ty of t h e Larderello region appears t o be r e l a t ed t o the emplacement of a post-orogenic batholith in to cont inenta l c rus t (Batini e t al., 198%; Puxeddu, 1989).

Natural Discharges

Hot water convecting in to t h e near-surface par t of a large hydrothermai sys tem may be dispersed by mixing with iateraliy flowing cold groundwater or discharged directly t o t h e surface. Only a minor amount of hea t energy is lost by conduction, but most is dispersed a s hot wa te r and vapor flows a t t h e surface. The processes af fect ing a deep f!uid penetra t ing t o t h e surface depend on a var ie ty of factors. Direct discharge depends on t h e availability of a suitable f r ac tu re system (or hydrothermal eruption vent, s e e below) and gives rise to a boiling spring, high in chloride and .mantied by silica sinter. Examples a r e the Champagne Pool, Waiotapu and the Pink and White Te r races of Rotomahana, New Zealand (Pla te 1.1; and s e e Henley et ai., 19841, Geysers a r e a special class of boiling discharge which have a periodic discharge due t o the geometry of t he conduit (Kieffer, 1984:. Often dilution precedes boiling of t h e mixed fluid a s i t finally moves t o t h e surface a s in t h e Ohaaki Pool a t Broadlands (Ohaaki) or t h e boiling springs of t h e Wairakei and Tauhara sys tems (Fig. I.4a).

Fluids which a r e diluted with respect t o the deep chloride wa te r form where in teract ion with near- surface aquifers occurs e i ther due t o high surface relief and groundwater flow or t o t h e proximity of t h e system margin. The natura l discharges of the Wairakei-Tauhara and Mokai sys tems a r e examples (Fig. 1.2).

Figure 1.3a shows schematicaliy the pressure distribution associated with various discharge phenomena. Drill-hoie da ta suggest t h a t pressure gradients in t h e deeper sys tem a r e generally about 10% above hydrosta t ic pressure with the excess pressure due t o t h e buoyancy of hot water re la t ive t o surrounding cold groundwater (Eider, 1966; Cathles, 1977; Grant et al., 19821, and in some cases (e.g,, Mokai) a demonstrable component of hydrostatic head due t o recharge f rom a reas of relatively high relief. An excess pressure gradient i s a requirement fo r i low through permeable media. Below a hot-spring vent, fluid expansion leads t o two-phase flow in the high- permeabili ty conduit. Phase separation may occur with the vapor discharging independently a t t he surface a s a fumarole o r in teract ing with groundwater t o produce a s team-heated water. As suggested i n Figure 1.3a, minor throt t l ing may occur aiong the flow path, but pressure drops a r e unlikely t o be greater than I bar. Where silicification isolates t h e conduit f rom the surrounding groundwater system, boiling, deep- sys tem iluid ex i t s t h e surface; but, where only partial isolation occurs through mineral deposition, the iiquid may itself i n t e rac t with surficial groundwater before reaching t h e surface a s a hot o r warm spring, In the

N N E z Si s - s e SSW L =

3 i l h l Y I , * .

0-2in E.0.r." Depaufi - 8 ,,,, ..,, ,",,.. d

Figure 1.2. DLstrikution of natural discharges in some active geothermal systems. T:?e field boundaries shown are based on the maximum resistivity gradient located by field silrveys reflecting the contrast between unmineralized groundwater and the ehlcride water present in the upper 500 meters of the geothermal systems. Fuma- roles, steaming ground, and outflows of steam-heated waters are indicated by the 0 syxilwl an6 hot-water discharges by the g symbol. The Location of the principal convective upflow for each field is indicated by the v . Notice that geothermal exyloration and production wells are situ- ated well away from natural features. Nu- merals designate features shown in the mixing diagrams of Figure 1.4.

Figure 1.2 (cont'd) a). Wairakei-Tauhara, New Zealand. These two fields are interconnected as shown by the resist-ivity and both show the occurrence of vapor discharge in the central region and hot..-wate qe on the margins followincj dilution. There is no evidence that water from take Taup penetrates either field, reaarge being derived from gr to the east and west. b). Waiotapu, New Zealand* This field has an extensive north-tosouth lateral flow originating in the vicinity of the 160,005 years hp. &cite domes to the north. Thermal features are related to major faults and a number of hydrothermal eruption craters have been recognized (major centers shown by the circles)--for full discussion see Fledempist and Wenley, 1985a. c). Mokai, New Zealand. Extensive lateral flow occurs from the vicinity of the caldera wall in the south toward the Waikato River to the north. Dilute hot springs occur north of the "field boundary" in the gorge of a stream following a major fault. d). Cross-section of the Mokai geothermal field showing the effect of lateral flow and dispersior. on the thermal structure of the system and distribution of natural features. (The cross section runs from the tap right-hand corner of Figure 1.2~ to the caldera wall south of well MK6).

Plate 1.2. Crater Lake, Ruapehu, New Zealand. Condensation of volcanic gas into the Crater Lake waters produces a fluid of pW 1.5 at about 55OC. The lake seldom overflows despite the presence of an incised channel (foreground) suggesting that much of the acid fluid drains through the core of the active andesite volcano producing an extensive high-level zone of advainced-argillic alteration. Interaction of this fluid with an underlying near-neutral pH hydrothermal system nay generate a gold depit of the Goldfield type [photo by permission, R B. Glover, DSIR).

PRESSURE 10 20 30 bars

Figure 1.3a. Pressure-depth relations in the upper portion of a geothermal system. The diagram shows the transition between the deep system pressure and the pressure within the high permeability fracture network or conduit below a hot spring. Below the hot springs, the pressure at a specified depth is due to the weight of a standing column of hot water; the pressure-depth relation is here designated ''hot hydrostatic". Deeper in a system pressures exceed hydrostatic so that flow is maintained through the permeable aquifer--this is shown as the ''hot hydrodynamic" curve. Some minor pressure discontinuities are shown to indicate the possible occurrence of minor throttles which may occur due to fracture geometry or silicification, but these are probably rare. Phase separation may occur resulting in the presence of fumaroles or (acid) steam-heated waters in the vicinity of a boiling hot spring (e.g., Norris Geyser Basin, Yellowstone, Champagne Pool, Waiotapu). The effect of raising or lowering the ambient groundwater piezometric surface may be gauged by redrawing the curve for cold-water hydrostatic pressure. For example, if the cold-water piezometric surface is at +20 meters and the hot-spring conduit is not isolated by mineral depi-

tion, dilution may occur near surface. Dilution occurs on the margin of a hydrothermal system due to the relative pressure of cold groundwater over that of the hot- water system.

example shown (Fig. 1.3a), deep mixing may occur where t h e pressure of cold water exceeds t h a t of t he hot upflow.

Exercise: The e f f e c t of relief, through a higher or lower piezornetric surface , may be gauged by adding cold water pressure curves t o Figure 1.3a corresponding t o higher and lower p iezometr ic surfaces. Try i t by drawing curves parallel t o t h e r e fe rence cold-water cu rve in t h e figure.

Haas (1971) has described t h e l imiting hydrosta t ic conditions f o r temperature a s a function of depth in hydrothermal systems. The l imiting condition (Fi.g. 1.3b) is t h e phase change t o vapor; liquid water rislng within a system boils a t t h e phase boundary with consequent format ion of a low-density vapor f rac t ion and a decrease in t empera tu re (for a disctission of reversible and irreversible boiling in hydrothermal systems, s e e Barton and Toulmin, 1961). As discussed above, hydrodynamic pressures prevail a t d e t h in g geothermal sys tems so t h a t a t , for example, 250 C the boiling-point depth is a t about 400 r a the r than 462 meters. The e f f e c t of salinity on t h e boiling point- depth relation is well known, but more recent ly t h e e f f e c t of gas pressure has been recognized (Sutton and McNabb, 1977) a s shown in Figure 1.3b. The l a t t e r e f f e c t makes i t particularly difficult t o obtain reliable depth information f rom es t ima tes of t empera tu re (e.g., f rom fluid inclusions) in fossil hydrothermal sys tems (Hedenquist and Henley, 1985b; Bodnar e t al., 1985, th is volume).

The distribution of springs re la t ive t o t h e geothermal system a s a whole is evident f rom t h e field maps shown in Figure 1.2. Areas occupied by hot- water discharge seldom represent more than about 5% of t h e a r e a of t he hydrothermal field itself. I t is also evident f rom these field examples t h a t t h e distribution of discharges i s strongly controlled by topography, t he presence of faults, etc.

In general, f ea tu res associated with vapor-flow from t h e deep system occupy higher ground. They range f rom fumaroles t o hot springs f e d by s team- hea ted surficial groundwaters t o s teaming ground which results f rom the boiling of s team-heated waters. The l a t t e r originate above two-phase ('boiling') zones in t h e deep convective system from which C 0 2 and H2S-rich vapor escapes, but a r e adsorbed in to surficial groundwater or condensate. Where H2S oxidation occurs due t o shallow interact ion with t h e atmosphere, low pH s team-heated wa te r s occur which a r e character ized by t h e presence of sul fa te and absence of significant chloride in solution a s well a s t h e lack of significant silica s in ter around the hot spring. C02-rich steam-heated waters, associa ted with i l l i t ic alteration, a r e also common marginal t o many fields (Mahon e t al., 1980; Hedenquist and Stewar t , 1985) t o depths of severa l hundred meters ; pH's a r e around 5 due t o dissolved C 0 2 and of ten result in ex t r eme corrosion of

T E M P E R A T U R E 100 200 300 O C

Figure 1.3b. Hydrostatic boiling-pint versus depth relations of hydrothermal fluids, showing the contrasting effects of salinity and gas content. As discussed in the text, observations from active systems suggest that pressure gradients a t depth are about 10% greater than hydrostatic allowing higher temperatures at shallower depth than shown here.

geothermal well casings. C 0 2 exsolution following mixing of hydrothermal fluid w ~ t h cool groundwater, and dissolution of the C O into groundwater may be t h e dominant process in t i e i r format ion r a the r than adiabat ic boiling.

Note: The t e rm "solfatara" encompasses s t eam - discharges such a s fumaroles, but i s now most commonly used t o refer t o volcanic g a d s t e a m discharges associated with sulfur deposition and advanced-argillic alteration. I t is o f t en used incorrectly in t h e discussion of me ta l t ranspor t in sub- s e a floor systems!

As noted above, acid sulfate-chloride waters derived by condensation of volcanic gas occur a t high levels in andesit ic t e r r ane and somet imes they may mix with meteor ic water and accumula te t o form c ra t e r lakes. Downward movement of such high-level wa te r s is of special in teres t and is now well known in explored geothermal fields in the Philippines and Taiwan. A t Ruapehu, New Zealand, for example, t he summit c ra t e r lake (Pla te 1.2) overflows

discontinuously despi te continuous input of volcanic vapor, me teo r i c water , and glacial melt . The c ra t e r lake itself has been shown t o be over 300 me te r s deep and is occupied by an acid sulfate-chloride water with pH = 1.25 a t 55OC. Presumably, t h e bulk of this acid fluid drains downward through t h e flanks of t he volcano causing advanced-argil l ic a l tera t ion e n route, and may encounter a normal hydrothermal system a t depth. Silica- and iron-enriched springs occur on the flank of the volcano a t lower elevations. These processes, a s noted, may be responsible for t h e acid- su l f a t e type o r e environments (see Hayba et al., 1985, this volume) and ra ises a l l so r t s of problems with respect t o terminology like hypogene and supergene!

Hydrothermal Eruption Vents

The hot springs described above a r e passive f ea tu res of t h e topography, but in some cases t h e system itself may gene ra t e high-permeability flow paths t o t h e surface. For example, a t Waiotapu (Fig. 1.2, P l a t e 1.1) t h e larges t single discharge of liquid f rom t h e deep system--the Champagne Pool--is independent of t he s t ra t igraphy and original topography and occupies t h e ven t of a hydrothermal eruption c ra t e r formed some 900 years b.p. Such eruption vents a r e now known t o have formed in a lmost a l l of t he New Zealand geothermal systems, but a r e less well known elsewhere due t o the f requent confusion of the erupt ive products with volcaniclastic breccias which may also be common in t h e vicinity. Hydrothermal eruption breccias a r e character ized by an absence of primary volcanic mater ia l and a r e generally polylithic and ma t r ix supported. Clas ts have a range of a l tera t ion s ty les and, together with s t ra t igraphic data , indicate an origin from depths up t o about 300 meters. An origin by gas exsolution has been proposed by Henley e t al. (1984) and by Hedenquist and Henley (1985a). Eruption breccias of shallower origin a r e also common in geothermal a reas and result f rom t h e in teract ion of vapor with surficial groundwater o r local removal of confining pressure, a s appears t o be t h e case for eruptions in Yellowstone (Muffler e t al., 1971).

H e a t and Mass Flow in Geothermal Systems

Table 1.2 shows hea t and mass output da t a from some geothermal systems. These a r e obtained by in tegra t ion of ground t empera tu re d a t a and physical measurements of t h e outflow r a t e s and temperatures of discharging hot springs and fumaroles. In some cases a n independent e s t i m a t e of t h e upflow is obtainable using measurements of t h e chloride content of river water up and downstream of a geothermal field (Ellis and Wilson, 1955; Fournier e t al., 1975). These d a t a may be r e l a t ed t o t h e convective upflow of high-temperature fluid in t h e sys tem assuming some knowledge of t h e upflow temperature . For example, a t Waiotapu the measured surface h a t flow is 600MW(h) (600 MW(h) = 600 x 10% Joules/s). Exploration drilling and geochemical d a t a sug e s t t h a t 8, t he fluid feeding t h e field is a t about 300 C. The enthalpy of s team-satura ted wa te r a t 3 0 0 ' ~ is about 1350 Joules/gm (see Henley and Brown, 1985, this

Table 1.2--Summary o f h e a t and mass f l ows i n some New Zealand geo the rma l f i e l d s

F i e l d T o t a l Heat Flow E q u i v a l e n t upf low

Wairakei Tauhara Waiotapu Ohaaki roadla lands) Mokai

volume, Fig. 2.3) so t h a t t h e mass flux of 300°C fluid is obtained by

Exercise: The Champagne Pool a t Waiotapu (Pla te 1.1) has a discharge of about 10 kg/s of 70°C water and, we es t imate , about 7 kg/s of s team. Calcula te t h e proportion of t h e to t a l convective upflow of hea t and mass in t h e system which is discharged by th is f ea tu re alone (the enthalpy of 70°C wa te r i s about 300 Joules/gm and of s t eam about 2600 Joules/gm).

Within t h e heat-flow budget t h e most difficult f ac to r t o assess is t he proportion of the upflow which may be dissipated by subsurface groundwater flow. On the basis of t he s ize of t he field and i t s deep temperatures , a t Broadlands (Ohaaki), t h e to ta l heat flow is thought t o be greater than 100 MW(h) and equivalent t o >75 kg/s of chloride water , but t h e observed surface hea t flow is less than a third of this es t imate . The principal outflow of hot wa te r f rom this field is t he Ohaaki Pool but, a t 10 kg/s, this accounts for only about 5 MW(h).

CHEMISTRY O F GEOTHERMAL DISCHARGES

Table 1.3 compares t h e chemis t ry of waters f rom natural f ea tu res with the chemis t ry of deep waters encountered by deep drilling. The principal discriminating f ea tu res with respect t o origin have already been noted, i.e., deep fluid character ized by CI>>SO and surficial s team-heated waters by sO~>>C!. In a given field, comparison of chloride contents in hot springs provides information about mixing processes and flow directions. Careful application of chemical geothermometer techniques may also provide some unique insights in to temperature pat terns in t h e underlying system and processes occurring during outflow (Fournier, 19811, and therefore provide an important guide for geothermal exploration.

Chemical relations between natura l discharges and the deeper chloride-water system a r e most commonly i l lus t ra ted by means of "mixing-diagrams." Figure 1.4 provides examples where t w o conservative quantit ies a r e compared; in this case chloride concentration and hea t content (enthalpy). The l a t t e r is frequently assumed t o be conservative during

Table 1.3--Summary of t h e chemis t ry of h o t s p r i n g s and 5eothermal f l u i d s

Concen t r a t ions i n mg/kg

F i e l d F e a t u r e t ° C pH(t ) C 1 SO4 H2S c02

W a i o ~ a p u Champagne Pool Well 7

Waiorapu Champagne Pool Well 80

Tauhara Crow's Nest Ka th l een S p r i n g Well 1

Mokai Nor the rn S p r i n g s Well 3

Tongonan B a n a t i S p r i n g Well 405

a Tauhara ., ,, ,,

1000 2000 CHLORl DE mg/ kg

I

1000 2000 CHLORl DE mg/kg

\ 2000.

E N T H A L P Y

kJ /kg

1000.

Figure 1.4a, b, and c. Fluid-mixing relations for the geothermal systems shown in Figure 1.2.

hydrothermal processes since conductive h e a t t ransfer i s a minor component of t he overall hea t budget and sys t ems a r e assumed t o b e in a s teady-s ta te with respect t o heat and mass t ransfer (heat re leased by a l t e ra t ion reactions is also a minor component). The principal processes occurr ing a r e dilution (mixing) and ad iaba t i c boiling during t h e irreversible expansion of t h e deeper system fluid a s i t rises and is subject t o less confining pressure. As discussed above, chemical t r ends due t o these processes a r e clearly shown in t h e d iagrams which then allow t h e in terpre ta t ion of t h e origin of individual t he rma l features. (In many cases, local mixing involves a s team- or conductively heated water wi th a t empera tu re in the vicinity of 150°c.) Some of t h e Wairakei-Tauhara springs, for example, show evidence of dilution prior t o boiling below a hot spring, while others, as at Mokai, a r e simply derived by dilution (Figs. 1.4a and c). The Champagne Pool a t Waiotapu is an example of t h e d i rec t discharge t o t h e surface of deep fluid (Fig. 1.4b).

EPITHERMAL ORE-FORMING SYSTEMS

Requirements fo r Ore Deposition

The transport chemis t ry of t h e epi thermal group of me ta l s has been reviewed by Barnes (19791, Weissberg e t al. (19791, and Henley and Brown (1985, this volume). In th is chapter , these d a t a a r e built in to t h e geothermal sys tem framework t o provide a n understanding of the origin of epi thermal mineral deposits in general based on t h e solution chemis t ry of t h e me ta l s and thei r response t o t h e t w o principal processes opera t ing in t h e upper levels of these systems--dilution and adiabat ic boiling. I t i s also impor t an t t o discuss t w o important in ter re la ted c r i t e r i a for o re deposition. These a r e (a) t he availabil i ty of me ta l s in solution and (b) t h e t ime required fo r t he o re depositing system t o opera te .

Figures 1.5a and b show t h e solubilities of gold, silver, and lead (representing the base meta ls) a s functions of t empera tu re and ligand concentration. As shown elsewhere (Ellis, 1970; Giggenbach, 1981; Henley e t al., 1984; Henley and Brown, 1985, th is volume), t h e pH of hydrothermal fluids in ac t ive and fossil sys tems is buffered by fluid + alumino-silicate react ions such a s t h e conversion of plagioclase t o mica and/or clay minerals. For low salinity fluids (CI = 1000 mg/kg), pH's a r e around 6.1 a t 2 5 0 ' ~ (i.e., on t h e alkaline side of neutra l pH) but about 1 pH unit more acid (5.1) fo r fluids an order of magnitude higher in salinity a t t h e s a m e temperature .

Fo r lead t h e dominant dissolution react ion in chloride solutions is

PbS + 2H+ + 2C1- = PbC12 + H2S

s o t h a t pH, chloride concentration and H2S content a r e the solubility controlling variables a t a given temperature . Clearly, t h e higher t h e salinity and lower t h e pH, the more me ta l i s dissolved, whereas high H S contents l imi t t h e solubility (Fig. 1.5a). (Note &a t in low chloride concentrations a t near

I I I 1 200 250 300

Tempemturn OC

log CI mglkg

1.0

0.1

B . - 0.0 1

. . I A g / : I i p b j I / : I ; : I / : I : i

9 OCEAN I I

2 3 4 5 Log CI q / k g

Figure 1.5. Gold, galena, and argentite solubility (mg/kg) versus (a) temperature and (b) ligand concentration for mineral-buffered hydrothermal fluids (calculated from the data summarizetj. in Henley et al., 1984, Chapter 9). In Figures 1.5a, b, and c, the fluid is considered buffered with respect to pH by the assemblage Kmica-Kfeldspar- quartz and with respect to fH2 by the empirical relation for the assemblage write- Fe-silicate-quartz derived by Giggenbach (1980). The pH of the fluid decreases to the right and f increases upward. Fig- ures 1.5a, b, ant c refer to examples discussed in the text. Eletal contents of the Broadlands system fluids (66 mg/kg H2S) are shown for reference. As an exercise, convert these data to values representing solutions containing 100 mg/kg H2S.

The slopes of the solubility curves shown relate to well-established thermodynamic data for the metal complexes considered, but relative solubilities may be in error due to the absence of reliable solubility constants for PbS and Ag2S., In the low-salinity fluids (see text), b~sulfide complexes of silver and hydroxy-carbonate complexes of lead may allow higher solubilities than calculated on the basis of chloride complexing alone.

The stippled region in Figure 1.5a emphasizes the temperature-salinity-metal concentration range of principal interest in epithermal studies.

Figure 1 . 5 ~ ~ for 250°c, includes an estimate of the solubility of Ag2S as Ag(HS)2- at low salinities and as chloride complexes at higher salinity. The ordinate in the case shows total reduced sulfur rather than H2S, and this introduces the curvature at low salinities (from Henley, 1986).

CHAPTER 1

Mil l ion o u n c e s A u + A g

1 0 0 0 0 .01 '7

C o r t e z , G o l d A c r e s b C a r l l n

A g : A U . -. -. O h n a k l P o o l

10 .- '---. -. - B R 2 2 S u l f i d e S c a l e k K u r o k o

u n n a l u a . 0

Figure 1.6. Au-Ag as a function of total gold and silver for a number of precious-metal deposits and the examples discussed in the text (modified from Graykal, 1981). Note that the ratio shown for case A is based on chloride complexing of silver, so that a more appropriate estimate for low-salinity fluids may be gauged from the ratio observed in the Broadlands BR 22 precipitate.

The host rocks of t he Broadlands system a r e really qui te unremarkable si l icic volcanic rocks and, a t depth, greywackes. Irrespective of t he host rock composition, a s shown above, significant gold transport and deposition can occur in relatively short t ime periods. The formation of an economic deposit is therefore more a function of the hydrology and chemistry of t h e system than i t is of t h e availability of unusual host-rock gold contents. Availability of t he metal(s) t o solution is, however, a f a c t o r which may contribute to the transport and deposition efficiency of the system a s a whole. In essence, fo r o r e exploration t h e recognition of the source of o re meta ls may pale in to insignificance re la t ive t o t h e recognition of a source for t he metal-transporting ligand.

A t Waiotapu, Hedenquist and Henley (1985a) showed, using appropriate e s t ima tes of t he pH, fH2, and m ~ 2 ~ of t h e deep fluid, t h a t t h e s ize of t h e system was such a s t o be f b l e t o supply a l l t he r e uired me ta l for a t leas t 10 years. In t h a t case in 8 10 years, with a constant hea t and mass flux equivalent t o the present, about 3.6 x lo7 grams (1.2 million ounces) of gold could have been transported, but since t h e hydrology of t h e near-surface sys tem appears t o allow only one s i t e with the focused

Initial solution:

e f f i c iency (%I

CI 1OOOmglkg Gold

million ounces

2 4 6

log t ime y e a r s

Figure 1.7. Gold deposition as a function of time, flow rate (10 kg/s) and process eff i- ciency for a fluid initially at 3 0 0 ~ ~ (see Fig. 1.5). The flow rate relates to that of a typical hot spring whereas the total flow of most geothermal systems is ten to forty times larger. How does this affect the position of the curves7 flaximum and minimum lifetimes for hydrothermal systems are shown for reference (for discussion, see Henley and Ellis, 1983).

depositional process, t h e overall efficiency of the hydrothermal system a s a m e t a l concentra tor is quite low, around lo%, the remaining 90% being disseminated widely through the shallow boiling zones.

Chemist ry of Systems Responsible fo r Ore Formation

I t is immediate ly obvious from Figure 1.5a, t h a t t h e prime requirement fo r t h e formation of a gold deposit is $ fluid relatively high in H2S but of low salinity ( 10 mg/kg C1). The systern temperature is of lower significance, a s shown by Figure 1.5b. By contras t , silver-rich base-metal deposits require fluids of high salinity ( seawater) and t h e e f f e c t of H2S concentration is secondary. The temperature coeff ic ients of solubility fo r t hese meta ls a r e large, but of less significance than salinity in determining me ta l t ranspor t capability.

In reviewing fluid-inclusion da ta from available studies on epi thermal systems, Hedenquist and Henley (1985b) confirmed the validity of these salinity cri teria. In many cases t h e low salinit ies of t h e fluids responsible for gold deposition were obscured by the presence in solution of dissolved gas, predominantly C 0 2 , which contr ibutes t o the f reezing point of t he

inclusion fluid in t he s ame way a s o the r solutes. For example, 4 wt.-% C 0 2 in solution depresses t he f reezing point by t h e s a m e amount , -1.7'~, a s 2.8 wt.-% NaCl. Since t h e molar C O /H S r a t i o of ac t ive geothermal-system fluids range2 f?om 10 t o L O O (Giggenbach, 1980), t h e high C O implies high H2S a s required by t h e solution model. h u e t o t h e scarc l ty of reliable gas analyses fo r fluid inclusions, t hc correlation of high- salinity, low-gas fluids with base metal-silver deposition has yet t o be fully demonstrated.

Although no s ta t i s t ica l analysis of fluid compositions is possible, i t appears f rom t h e d a t a available f rom well-explored geothermal sys tems t h a t salinit ies r e f l ec t host rock and crus ta l sett ing. This is shown schemat ica l ly in Figure 1.8, which makes t h e hoc assumption of normal f requency distributions fo r - fluid composit ions in d i f ferent crus ta l environments. Salinit ies fo r basalt-hosted sys tems a r e lower t han those fo r sys tems hosted by si l icic volcanic and these in turn a r e lower t han fo r andesite-hosted systems. Doesn't th is broadly r e f l ec t t h e ore-host relationships seen in many districts? Fluid-inclusion d a t a suggest t h a t silver deposits of t h e Creede- type formed f rom fluids of even higher salinity. As discussed e lsewhere (Hedenquist and Henley, 1985b), such high-salinity fluids a r e encountered in some t e r r e s t r i a l geothermal systems. Such fluids, some with ex t r eme ly high salinity, occur in sys tems typified by those of t h e Imperial Valley, California, within which evapor i tes a r e present (Rex, 1983), ref lec t ing both t h e t ec ton ic

se t t ing , c rus t a l r if ts , and ambien t c l imate . Using regional geologic da t a , therefore , i t m a y b e possible t o discriminate hydrothermal sys tems, both anc i en t and modern, which could hos t si lver-base me ta l mineralization f rom those potent ia l ly hosting gold, if the i r gas con ten t s were high enough.

Chemical and Physical Processes in Ore Format ion

If i t is a c c e p t e d t h a t present-day ac t ive geothermal sys t ems a r e t h e a r che types of those responsible fo r ep i the rma l o r e deposit ion in t h e past , t h e discussion of t h e physics of na tu ra l geothermal discharges (above) becomes immedia te ly re levant t o t h e discussion of ore- forming processes. T h e natura l discharge of t hese l a rge geo the rma l sys t ems i s focused on highly localized f e a t u r e s such a s hot springs whose locations r e f l ec t both topography and underlying geological s t ruc ture . As shown above, t hese f e t u r s 4 5 a r e usually conf ined t o a n a r e a of t he o rde r 10 m , less than 5% of e s u r f a c e a r e a of t h e pa ren t sys tem

7 Y (say 7.5 x 10 m ). The f e a t u r e common t o a l l these discharge paths was t h e progressive pressure drop f rom t h a t of t h e upflow sys t em t o t h e ambien t pressure of near-surface aquifers or t o a tmospher ic pressure. Boiling and dilution by near-surface wa te r s a r e t h e accompanying processes which lead t o mineral deposition a long these flow paths.

Although t h e solubil i ty and solution chemis t ry of a f ew m e t a l s have been out l ined exper imenta l ly t o an e x t e n t suff ic ient t h a t t he i r gross t ranspor t in these

I GEOTHERMAL FLUID COMPOSITIONS I Change in equivalent - wt % N o ~ l due to add~t~on of C02

I VOLCANIC HOSTED SYSTEMS I EVOLVED CONTINENTAL BRINES

recharged systems Salton Sea 25 wt % Cheleken 26 wt % Cesano 13 wt %

EPITHERMAL FLUID COMPOSITIONS ~ Correction to average equ~valent wt % N a ~ l for C02

1 Gold-silver ores Kuroko ores Base rnetol ores

- NaCl wt %

Figure 1.8. Distribution of fluid salinities in the earth's crust in relation to host-rock and crustal environment. A normal frequency distribution has been assumed for each fluid type in the absence of evidence for a continuity of compositions. For discussion, see text and Hedenquist and Henley, 1985b.

hydrothermal sys tems may be mapped, t h e understanding of ore-deposi t~onal mechanisms is f a r less advanced; indeed, n o exper imenta l studies have been a t t e m p t e d in view of inherent difficulties. Simple k inet ics suggest, however, t ha t optimum deposition r a t e s a r e achieved from the most supersa tura ted solutions. This allows t h a t discussion of o re deposition rnay concentra te on t h e major changes in solution chemis t ry consequent on boiling or dilution.

The e f f e c t s of adiabat ic boiling and dilution on solution chemis t ry have been discussed elsewhere (Ellis, 1970; Henley and Brown, 1985, this volume; Henley e t al., 1984; Drurnmond and Ohmoto, 1985). Boiling is an especially important process because the format ion of only a f ew percent of vapor allows t h e loss of more than 90% of dissolved C 0 2 , with a concomitant pH increase by more than one pH unit and t h e loss of H2S.

Consider t h e e f f e c t s of boiling and dilution on t h e case history solutions A and B discussed above. In both cases t h e t empera tu re change due t o these processes is taken for i l lustration a s 50°C. Lead and silver supersaturations, a t t a ined by dilution with f resh water (with host-rock pH buffering), a r e 500 and 100, respectively, due t o changes in temperature , pH, and chloride concentration.

A t Creede , Colorado, fo r example, Hayba (1984) and Hayba et al. (1985, th is volume have shown t h a t t t he mixing of high-salinity (7.2 x 10 mg C1-/kg) fluid with surficial steam-heated water was contemporaneous with o r e deposition. The exsolution of C 0 2 during such mixing cooling may also contr ibute t o base-metal and silver deposition in some environments. The supersaturations a t ta ined due t o adiabat ic boiling (~pH=+1 .5 , AH+-80%) a r e 2000 and 300 for P b and Ag (chloride spec~es ) , respectively.

Figure 1.5a suggests t h a t t h e mineral-buffered solubility of gold increases with dilution. This occurs in response t o t h e increasing pH of the sil icate- buffered solution a s t empera tu re falls and clearly gold deposition cannot be re la ted t o simple dilution. If ac idic fluids a r e t h e dilutant, a s may be t h e case in t h e high-level andesite-hosted systems, fluid mixing can cause gold deposition. In this case, a re la t ive pH decrease lowers the solubility of gold allowing a paragenesis of gold-kaolinite-alunite t o occur in t h e mixing zone between high-level acid-sulfate-chloride wa te r s (with associated advanced-argillic a l tera t ion) and deeper near-neutral pH chloride waters (with associated propylitic alteration). The high arsenic content of t he o r e may also r e f l ec t this environment (see below).

In many epi thermal deposits, gold is clearly associa ted with mineralogical indicators of boiling (open-space filling calc i te , adularia, sulfides). The fluid da ta f rom Broadlands (Brown, 1985) also very clearly demonstra te the ef fect iveness of boiling a s a process for t h e deposition of gold, silver, and copper, e a c h in solution a s a bisulfide complex (Henley and Brown, 1985, th is volume). In this case, t h e init ial pH increase due t o C 0 2 loss may undersatura te t h e fluid but sustained loss of t he more soluble H2S with t h e format ion of only a few percent s t eam leads t o supersaturation and deposition.

The loss of H2 due t o boiling gives an apparent increase in t h e redox s t a t e of t h e fluid ( re la t ive t o t h e pyrite-pyrrhotite stabil i ty boundary) and also leads t o a n increase in solubility inside t h e H2S stabili ty field. Quant i ta t ive gold deposition may be expec ted if t he final redox s t a t e corresponds t o some point in t h e SO; stabil i ty field but this poses some severe headaches with respect t o t h e stabil i t ies of accessory minerals and t h e kinetics of H S H20-SO: reactions. (For fu r the r discussion, s e e fIinley and Brown, 1985, th is volume, and Reed and Spycher, 1985, this volume). As discussed by Thorstenson (19841, modelling of t h e redox response of a fluid subject t o such a non- equilibrium process a s boiling is f raught with difficulty both in concept and in dealing with the redistribution of e lec t rons over the large number of e l ec t roac t ive couples available in natura l fluids.

In the discharge of well BR22 a t Broadlands, gold- and silver-ore deposition is largely comple te within a few seconds of boiling a t t h e well-head throt t l ing pla te (Brown, 1985). This observation suggests t ha t t h e loss of ligands (e.g., HS- a s H2S) is most important. React ion equations like

i l lus t ra te this and emphasize t h a t in discussing t h e deposition of gold under non-equilibrium conditions i t i s necessary t o identify sources of e lec t rons fo r t h e fu r the r reduction of aurous ions t o the metal.

Wgh gold a t concentration of t h e order of Ipg/kg (3 x 10- molal), t he re a r e a number of possibilities because of t h e ext remely low elect ronegat iv i ty of gold. The oxidation of o ther dissolved meta ls and of reduced sulfur a r e all possible sources. The deposition of silver in e lec t rum proceeds in the same manner.

A role for arsenic in gold t ranspor t and deposition has of ten been suspected. Some exper imenta l d a t a suggest t h a t arsenic complexing may increase the solubility of gold in reduced sulfur solutions, but t he field d a t a from Broadlands compar ing observed gold content with the solubility ca lcula ted from bisulfide complexes suggest t h a t t h i s e f f e c t is minor. The high solubilities of arsenic sulfides in alkaline-sulfide solutions (Seward, 1984) suggest t h a t thioarsenide complexes occur, but t hese may well be supplanted by arseni tes in moderate19 acid-neutral pH deep, system waters. If so, thioarsenides may provide a sink for H2S in residual, relatively high-pH wa te r s derived by boding. Arsenic concentra t ions in geothermal waters a r e usually in t h e range 1-10 mg/kg and could adequately lock up residual H S.

~ i e l c f exper iments (Brown e t al., 1983) involving geothermal waters suggest t h a t amorphous arsenic sulfide may be precipitated by acidification through react ions of t h e form

Natural examples of this process occur in t h e Tamagawa Hot Springs, Japan (Nakagawa, 1971) and in

a number of springs in the New Zealand and Yellowstone geothermal areas. A t Champagne Pool, Waiotapu, for example, t h e large surface a r e a results in hea t loss t o t h e a tmosphere and internal convection t o a depth of about 7 5 meters. With a mass influx of about 20 kg/s, deep fluid enter ing the pool i s quickly cooled and t h e pH buffered internally t o about 5 by C02-HCO; equilibration (Hedenquist and Henley, 1985a). A similar internally buffered process could be invoked fo r deposits like t h e sediment-hosted deposits a t Getchell , Nevada.

Both in natura l occurrences in New Zealand and in t h e geothermal field experiments, t h e amorphous arsenic sulfide precipi ta te is ore grade in gold and silver. Seward (D.S.I.R., personal communication, 1985) and Ellis (1969) have suggested t h a t colloidal arsenic sulfide scavenges gold from solution. Recrys ta l l iza t ion t o realgar and orpiment with f r e e gold and associa ted minerals such a s s t ibni te and cinnabar may then account for t he late-stage, low- grade o re assemblages occurring a t Getchel l and elsewhere. Together with t h e zonation of t r a c e me ta l s observed in t h e upper pa r t of t he Broadlands and other geothermal sys tems (Ewers and Keays, 19771, t h e association of gold with arsenic is commonly indicative of shallow depositional environments.

As noted above, t he e lementa l association of gold and arsenic is also common in t h e Goldfield-type deposits, e.g., Goldfield, Nevada; Summitville, Colorado. In these environments mixing of descending high-level acid-sulfate waters with normal hydrothermal chloride-dominant waters may account for t h e deposition of gold and arsenic (as enargite) in association with an advanced-argillic a l tera t ion assemblage, t h e me ta l s derived from the deep system.

The association of gold with organic m a t t e r has raised questions concerning gold transport and deposition in deposits such a s Carlin and others in northern Nevada. Interaction of normal hydrothermal- system waters with organic-rich calcareous host rocks leads t o the rma l matura t ion of the hydrocarbon a s suggested by Ilchik (1984) for t he Alligator Ridge deposit and documented fo r the Cer ro P r i e to geothermal sys tem (Barker and Elders, 1979). L i t t l e e lse distinguishes the possible mineralogical response of t hese rocks t o hydrothermal a l tera t ion and i t is most likely t h a t t h e Nevada deposits result from sub- hot spring boiling a s discussed above. Finely dispersed carbon may have a secondary role t o play in t h e ext ract ion of gold from solution, although even th is in teract ion is not well supported by mineralogical d a t a (Wells and Mullens, 1973).

Host-Rock React ions

Whereas mixing and boiling a r e processes common t o a l l hydrothermal systems, specific in teract ions with host rocks which may deposit o re a r e f a r less common. An exception is t h e in teract ion of relatively low-pH, high-salinity fluids with carbonate rocks which may, through loss of acidity, deposit massive replacement ores of silver and base me ta l s (e.g., Taxco, Mexico). Interaction of a high-H2S fluid with a n iron-rich sediment may be a r a re possibility for t h e format ion of a gold-pyrite assemblage.

SUMMARY

The original question posed a t t he beginning of this chapter was "why study geothermal sys tems in t h e con tex t of t h e origin of epi thermal o re deposits?" I t is c lear f rom t h e foregoing paragraphs t h a t such s tudies highlight a number of important f ac to r s necessary for t h e understanding of chemical and physical processes in epi thermal systems. These may b e summarized a s follows:

1. The ac t ive high-temperature geothermal sys tems in volcanic-rock t e r r anes a r e t h e archetypes of those sys tems responsible for epi thermal precious- and base-metal o re deposits in analogous ancient terranes. Porphyry copper-molybdenum deposits represent crust-magma in teract ions a t somewhat deeper levels within calc-alkalic volcanic-rock hosted sys tems and Kuroko-type massive sulfides a r e essentially sea-floor telescoped equivalents of the t e r r e s t r i a l epi thermal deposits.

2. Studies of ac t ive geothermal sys tems provide insight in to the physical processes governing flow t o surface discharge f ea tu res such a s hot springs and near-surface in teract ion with ground and other hot waters. Geological s t ruc tu re provides a principal control, but in many geothermal fields hydrothermal eruptions focused a t depths of about 300 me te r s provide focused flow paths t o t h e surface. The cha rac te r i s t i c hydrothermal eruption breccias formed by these events have been recognized in a number of epi thermal precious-metal deposits such a s Round Mountain, Nevada (Tingley and Berger, 1985), and McLaughlin, California. Figure 1.9 provides an in teres t ing case history for t h e reconstruction of a n epi thermal system based oq field observations, fluid inclusions, and isotope techniques. Such reconstructions a r e important t o t h e exploration geologist for target ing drilling and for comparison with o the r d is t r ic ts under exploration where complementary but piece-meal d a t a may be available.

3. Studies of ac t ive geothermal sys tems provide insight in to t h e range of fluid compositions present in t h e crus t . In conjunction with fluid and laboratory studies i t may be shown t h a t t he format ion of gold- silver (base-metal poor) epithermal deposits requires fluids of low salinity and high gas content (i.e., H2S in association with C O ), whereas silver-rich base-metal deposits (relatively Tow in gold) require high-salinity, low-gas fluids in their respect ive hydrothermal systems.

The salinit ies of hydrothermal fluids r e f l ec t thei r volcanic-tectonic environment. For example, andesite-hosted hydrothermal sys tems a re , perhaps through high-level in teract ion with volcanic gases or t h e involvement of deep, connate waters, more saline than those of silicic volcanic-rock ter ranes , but less saline than those encountered where evapor i te sequences occur (grabens, some caldera moats); t h e l a t t e r ref lec t ing also a c l imat ic control. Although acid-sulfate waters a r e locally encountered in silicic- rock hosted systems, they a r e quite abundant in high- level andesitic-rock terranes, and may be important components in t h e formation of cer ta in types of gold deposits.

Such studies also highlight t h e important role of

ROUND MOUNTAIN, NEVADA GEOLOGY $ ALTERATION < r Y m T n ( i i s r L Bsi.ei 1184

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ROUND MOUNTAIN, NEVADA HYDROLOGIC RECONSTRUCTION I * "."ley

.. . .- .- - -- - m e r e .

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6OOrneti.l 100 JOB 200 150 D

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Figure 1.9a, b, and c. Distribution of alteration types and fluid-inclusion data for the Round Mountain gold deposit, Nevada. Data from Tingley and Berger, 1985. The topographic outline shown is drawn from the viewpoint shown in a).

Hypothetical reconstruction of the Round Mountain hydrothermal system (25 million years b.p.) emphasizes the dissection of the system by erosion. Much controversy has centered on the so-called hot-spring origin of this deposit; clearly the deposit is hot-spring related but is not itself a hot spring1 There seems to be little advantage to the use of the term "hot- spring-type deposit" over the generally applicable "epithermal" unless very specific evidence is available for a given deposit relating it to a hot-spring environment (e.g., silica sinter, eruption breccia) as at McLaughlin, California. The hot-spring terminology may indeed be far tcx, restrictive for the guidance of exploration programs even though the recognition of hot- spring features is a very important tool for the exploration geologist.

gas composition in controlling t h e physical processes within sys tems and their me ta l t ranspor t capability, a s well a s demonstra t ing t h a t two processes (adiabatic boiling and mixing) provide the principal controls on fluid chemis t ry in the epi thermal environment. Access t o some gas source ( C 0 2 + H S) becomes, therefore , paramount in determining whe&er a given low-salinity system may or may not develop a gold deposit. As discussed elsewhere (Henley, 19861, t h e r e may be t ec ton ic controls here (e.g., access t o gas f rom subduction zone metamorphism?) which introduce t h e link between t ec ton ic sett ing, fluid chemistry, and o re formation, which together provide guides t o o re search.

4. Active geothermal systelns provide the opportunity for studying t h e deposition of t r a c e me ta l s such a s gold and t h e init ial concentra t ions of ore- forming e lements in the deep system. Geothermal systems such a s Broadlands show t h a t me ta l concentrations may b e regarded a s close t o sa tura t ion and demonstra te t h a t host rocks play only a passive ro le in determining whether sys tems a r e likely t o deposit o re near t h e surface. The principal control on t h e l a t t e r is t he provision of permeable f ea tu res within which flow may be focused and adiabat ic boiling and/or dilution may occur. The kinetics of me ta l deposition during irreversible boiling a r e not well understood and hard t o model realist ically, but pH and H S concentration appear t o be more important than reTatively slow redox reactions. More specialized environments involving gas buffering of pH o r mixing with acid waters may be indicated by t h e association of gold with orpiment-realgar or enargite-alunite assemblages.

EPILOGUE

If physical and chemical processes alone a r e t h e control on whether or not a system may deposit an o r e body, what chance does t h e exploration geologist have in target ing exploration effectively? The answer is in th ree parts: (a) Since act ive geothermal sys tems a r e relatively similar in t e rms of the distribution of a l tera t ion minerals, these assemblages may be used t o provide key d a t a on t h e level of exposure within a system; clays and chlorites, for example, a r e qui te sensit ive indicators of temperature 'of format ion (Browne, 1978; Henley and Ellis, 1983). Recognition of o ther depth indicators such a s hydrothermal eruption or vent breccias is also a powerful target ing tool. Similarly t h e recognition of t race-e lement enr ichments (As, Sb, Hg, TI) provides a clue t o s t ructura l level in t h e system. (b) Good qlold-fashioned' s t ructura l s tudies have been much neglected in r ecen t decades, but they probably s t i l l provide t h e most e f f ec t ive means of target ing drilling since s t ructure , particularly f r ac tu re analysis provides t h e main clue t o flow s t ruc tu re in t h e fossil system.

Knowledge of t h e hydrodynamic character is t ics common t o ac t ive geothermal sys tems also provides a guide t o t h e exploration of fossil systems. Additional f ac to r s t o remember a r e t h e e f f e c t s of contemporary

or subsequent volcanism and tec tonism which obscure and dismember t h e original s t ructure . A useful exercise (or, in modern parlance, !'thought experiment") is t o imaginatively modify copies of Figure 1.1 by inclusion of layers of volcanic mater ia l and/or f au l t dissection t o explore some of the problems which ar ise in reconst ruct ion based on geological evidence. In tackling th is exercise, remember the three-dimensional aspects of t he systems. (c) Discriminant analysis i s a t e r m used in s ta t i s t ics t o describe a method for classifying mul t ivar ia te observations in to groups. In t h e above paragraphs, the variability of surface expression and chemis t ry has been s t ressed but a s noted by Koch and Link (1970, p.326) this is ra ther pointless if no ways a r e established t o so r t o u t t hese kinds of variability. This is t h e essence of t h e design of sc ient i f ic experiments, but exploration for epi thermal o re deposits i s seldom e f fec t ed in th is manner, i s i t? Since t h e coincidence of a suitable flow s t ruc tu re , f r a c t u r e pat tern and system chemistry is responsible for ore-metal deposition ra ther than dispersion, t h e probability of exploration success i s t h e f a c t o r requiring ear ly determination in a n exploration program. Fac to r s such a s a previous mining history a r e obvious high-score fac tors a s i s recognition of s t ruc tu ra l style. Modern techniques such a s isotope analysis and fluid-inclusion studies should also be used in a discriminatory manner; e.g., fluid inclusions may very ear ly in a project determine whether a sys tem is relatively dilute and therefore not likely t o host base metal-silver o re or relatively gassy and the re fo re quite likely t o have transported and perhaps deposited gold. Understanding of the geochemistry of me ta l t ranspor t and deposition rnay be used in t h e s a m e way.

ACKNOWLEDGMENTS

Discussions with K. L. Brown, R. G. Allis, P. B. Barton, and P. M. Bethke have been valuable in preparing this overview. The musical accompaniment of J a m e s Galway was especially soothing t o t h e writer and may well provide solace t o t h e reader!

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Chapter 2 A PRACTICAL GUIDE TO THE THERMODYNAMICS OF GEOTHERMAL

FLUIDS AND HYDROTHERMAL ORE DEPOSITS R. W. Henley and K. L. Brown

INTRODUCTION

In trying t o understand t h e depositional processes which led t o o re deposition in fossil hydrothermal systems, we a t t e m p t t o reconst ruct t he chemis t ry of t h e fluid phase f rom observation of i t s relics (e.g., a l tera t ion minerals, fluid inclusions). We may also a t t e m p t t o thermodynamically model t h e chemical changes experienced by this fluid a s i t passes upward through a vein, vents t o t h e seafloor, boils o r mixes with other waters, e tc . A number of important assumptions a r e made; one is t h e assumption of equilibrium and another is t h a t t h e thermodynamic d a t a base is sound.

Analyses of fluids discharged from geothermal wells, together with drill-core data , allow t h e opportunity t o independently check the validity of t he thermodynamic d a t a base and t o observe directly, chemical processes leading t o t h e deposition of gold, base-metal sulfides and common gangue minerals like quar tz and calcite. The calculations involved a r e not trivial, but a r e essential t o t h e understanding of epi thermal or any other type of hydrothermal o re deposit.

To i l lus t ra te these procedures, we shall examine t h e discharge of one production well in t h e Broadlands geothermal field in New Zealand. Through t h e use of thermodynamics, t h e amount of information we shall re t r ieve about t h e reservoir and depositional processes i s qui te astonishing. We shall then turn t o some review questions t o consider implications fo r t h e format ion of some hydrothermal o re deposits.

In this chapter we have t r ied t o follow a pragmat ic course, avoiding the temptat ion t o overindulge in the (essential) nuances of thermochemistry a t t h e expense of t h e proscribed goal--a practical understanding of hydrothermal processes. A wider discussion of geothermal chemistry, a broader d a t a base, and references t o t h e current l i te ra ture a r e presented in Henley et al. (1984).

GEOLOGICAL CHARACTERISTICS O F THE BROADLANDS GEOTHERMAL SYSTEM

The geology and a l tera t ion mineralogy of t h e Broadlands field have been recently reviewed by Weissberg e t al. (1979). Figure 2.1 shows t h e location of exploration and production wells in t h e Broadlands field and Figure 2.2 shows in cross-section t h e s t ructure and stratigraphy of t h e field in relation t o measured subsurface temperatures. The host rocks for t he hydrothermal system a r e mostly si l icic ash-flow

originally containing qua r t z and andesine with a glassy or fine grained groundmass and minor hornblende, bioti te, magnet i te and other common accessories. The geology and a l tera t ion mineralogy of cores from several wells a r e summarized in Table 2.1. As well a s ubiquitous pyr i te and lesser pyrrhotite, copper, lead and z inc sulfides have been recognized in veins and vugs in core frorn some wells, and a r e commonly associated with adularia and calc i te . A t t h e surface in the Ohaaki Pool, i n t e rmi t t en t deposition of an amorphous sulfide precipi ta te has oc'curred on the silica sinter formed a t t h e edge of t h e pool. This orange precipi ta te is o re g rade with r e spec t t o gold and silver and rich in As, Sb, TI, and Hg. A similar

0 , I B A Y O F

P L E N T Y

B " 0 ~ O L A N O S

.A?: . ' "PO b 0 1 0

27 t P, Ctj T H E R M A L

A R E A

1 500 l000rn - Figure 2.1. Location of geothermal wells and

hot springs at Broadlands New Zealand. Wells marked, 8 , intersected base-metal sulfide minerals at depth; wells marked, 8 , deposited precious-metal-bearing, antimony-rich precipitates at the surf ace; and wells marked, , have both surface precipitates and base-metal sulfides at depth (reproduced with permission from Weissberg et al., 1977).

Table 2.1--Distribution of base-metal sulfides in Broadlands drill holes (reproduced from Weissberg et al., 1979)

Base-metal sulfide zone

Depth Depth Temperature Rock types in Associated Well of well range (m) range (OC) base-metal hydrothermal No. (m) sulfide zone minerals

Minimum Maximum Minimum Maximum

rhyolite tuff, greywacke

quartz, adularia, albite, calcite, pyrite, illite, chloriLe

conglomerate, greywacke

quartz, adularia, chlorite, illite

quartz, adularia, chlorite, illite, pyrite

argillite, conglomerate

ignimbrite, tuff

calcite, quartz, chlorite, adularia, albite, pyrite

quartz, illite, chlorite, calcite, pyrite, pyrrhotite

Broadlands Dacite, Rautawiri Breccia, greywacke

Upper Waiora pumiceous tuff breccia

calcite, illite, wairakite, quartz, pyrite

tuff quartz, pyrite, sphene, zoisite

tuff, tuf faceous sands tone

illite, pyrite, chlorite, sphene

precipi ta te formed in t h e discharge disposal channels of a number of exploration wells. The depth zonation indicated by these observations is also ref lec ted in chemical analyses of sulfide separa tes (Ewers and Keays, 1977) and recal ls t h e general pat tern of me ta l zonation encountered in many epi thermal o re deposits.

Exercise 1. For la ter convenience, complete the conversion of t h e liquid sample analysis t o molal units. (n.b. Cindicates analytical to ta ls for dissociated species; e.g.

In order t o examine chemical relationships between the reservoir mineral assemblage and the fluid phage, t h e da ta must f i rs t be recombined t o t h e single high-temperature and pressure phase which f ed t h e well. For this purpose a hea t balance equation is used (with t h e assumption t h a t negligible hea t loss occurs) t o ca lcula te t h e mass of liquid water converted t o vapor a s t h e fluid rises, a t several m/sec to the wellhead.

FLUID CHEMISTRY

Table 2.2 contains analyses of t h e liquid water and s t eam separa ted f rom the discharge of BR22 a t Broadlands. The discharge from this well is fairly typical, although higher temperature zones a r e encountered a t depth elsewhere in t h e field.

R. W. HENLEY & K. L. BROWN

Table 2.2--Analytical data for fluids discharged from well BR22, Broadlands, New Zealand

Liquid sample

Collection Pressure pH20 Na K Ca Mg C1 B SO4 Sio2 CNH3 XHC03

11 b.a. 7.39 870 188 1.2 0.1 1432 42 2 705 6.1 216

11.75 3 10 x moles kg-'

Steam sample

Collection Pressure C02 H2S NH3 H2

6.6 5.0 lo3 x moles kg-'

Figure 2.2. Geologic cross-section of the Broadlands, New Zealand, geothermal field (reproduced with permission from Weissberg et al., 1977) .

where H is t h e enthalpy of t h e to t a l discharge (T.D.), and of liquid water (1) and vapor (v) a t t he separation pressure; y is t h e s t eam fract ion a t t he sampling pressure. Note t h a t 100 y is t h e percent of s t eam by mass in t h e two-phase mixture.]

HTD is determined by physical measurement and for this well is 1160 (1.20) J/g. Substi tuting appropr ia te values of HI and Hv from tables of thermodynamic d a t a for water (see Henley et al., 1984, Appendix 1111, equation (1) gives y = 0.19 .

Digression No. 1

For convenience in this chapter , Figure 2.3 shows s team fract ions for pure water a s a function of enthalpy, pressure and temperature . To use this diagram loca te t h e discharge enthalpy on t h e two- phase coexistence curve--this point represents conditions in t h e reservoir. Then project downward a t constant enthalpy t o the sampling pressure and read off t h e s t eam fract ion present under the reduced pressure conditions. This procedure, or more precisely t h e hea t balance equation, is t h e basis for t h e calculation of changes in chemistry and t empera tu re during boiling (adiabatic, closed system) of hydrothermal fluids in an ore zone (for example, s e e Fournier (1985, this volume) and Reed and Spycher (1985, this volume)). Hea t and mass balance equations a r e also t h e basis for similar calculations involving the mixing of fluids.

A mass balance equation is required to recombine t h e analytical data. For a component, X

XTD = (1 - y)Xl + yXv ( 2 )

so tha t , fo r example

S i 0 2 TD = 0.81 x 705 mg/kg

and converting t o molality by dividing by t h e molecular weight of S i 0 2

Si02,TD = 9.52 x moiesikg

where a volati le component such a s C 0 2 (534.3 x 10- 3

moleslkg in t h e vapor in this example) is considered

''~,TD

= 0.81 X ~ H C O + 0.19 X 534.3 X moleslkg 3

Figure 2.3. Enthalpy-temperature- pressure-density relations for water to 5 0 0 ~ ~ and 600 bars (from Henley et al., 1984).

ENTHALPY ABOVE 0' C, Joules/grn

Notice t h a t for C 0 2 we combine a l l t h e analyt ica l ca rbona te carbon (i.e., C 0 2 , HCO3, C 0 5 , and C 0 2 ), whereas fo r si l ica we may reasonably assume t h a t tke sil ica content of t h e s t eam is negligible.

Exerc ise 2. Comple t e Table 2.3 t o obta in t h e t o t a l discharge composit ion of t h e well.

FLUID-MINERAL EQUILIBRIA: ALTERATION MINERALOGY

What controls t h e concentra t ion of e a c h component represented in t h e analysis?

Fi rs t we examine silica. I t is generally assumed t h a t since qua r t z is ubiquitous in t h e reservoir , i t s solubility controls m . 0 2 in t h e fluid. We can check th is through availazfe exper imenta l d a t a fo r t h e solubility of qua r t z in water (reaction 3)

Figure 2.4 shows original solubility d a t a fo r q u a r t z (Kennedy, 1950; Morey et al., 1962). The observed concentra t ion of silica in t he aquifer fluid is 570 mg/kg (Table 2.3) so t h a t inspection of Figure 2.4 yields an aquifer t empera tu re of 2 6 5 ' ~ .

Exerc ise 3. Now check t h e measured enthalpy of t h e discharge against s t eam t ab l e s or Figure 2.3 t o find a t wha t t e m p e r a t u r e i t corresponds t o t h e enthalpy of vapor- s a tu ra t ed liquid. You have probably guessed t h e answer anyway; t = 2 6 5 ' ~ .

C a n you now infer t h a t msio2 is controlled by t h e solubility of quar tz?

Equilibrium constants fo r th is and other useful reac t ions a r e given a s t empera tu re dependent regression equations in Appendix Table 2.AI. These equations apply over l imi ted t empera tu re ranges. For example , for quar tz , t h e equat ion given assumes a l inear relationship between log K and 1/T f rom 200 t o 2 8 0 ' ~ and conforms t o t h e solubility d a t a recommended by Fournier and P o t t e r (1982). Various o ther equations have been published in t h e pas t and r e l a t e t o t h e e x t e n t of exper imenta l d a t a available, t h e authors ' bias, t h e regression technique used, and t h e assessment of errors. In th is case , t h e equation yields a slightly higher aquifer t e m p e r a t u r e which may or may no t be real. The enthalpy measurement and exper ience in a la rge number of wells confirm t h a t qua r t z cont ro ls t h e si l ica con ten t of t h e aquifer fluid above about 200°C (see Fournier, 1985, this volume).

Other components m a y be examined in a similar way. For example, fo r t h e hydrolysis reac t ion

3KA1Si3O8 + 2 w Kspar

which is here wri t ten with conservation of aluminum in t h e solid phases.

Appendix Table 2.A1 gives log K265 = 7.897 a t 2 6 5 ' ~

R. W. HENLEY & K. L. BROWN

Table 2.3--Total d i s c h a r g e composi t ion f o r w e l l BR22

moles kg-' x l o 3

Na K Ca Mg C 1 B Si02 C N H 3 Z C O 2 H2S H2 CH4 N 2

and log K = 21og a ~ + + 2 pH (5)

where a is the thermodynamically e f f ec t ive concentration of, in this case, potassium ion.

Digression No. 2

We need for a moment t o consider ac t iv i t ies , ai , and act iv i ty coefficients, Y i , for solution species designated i. Yi provides t h e link between t h e rea l messy world and tha t of ideal thermodynamics.

For t h e silica problem, we adopted Si(OH)4 = I; but this is allowable only for neutrally charged species in relatively dilute solutions up t o say 1 molal. Individual ion activity coefficients a r e tr icky t o calculate, but a r e usually approximated in dilute solutions up t o say 3 molal using the extended Debye-Hiickel equation. For t h e high-temperature calculations discussed here, t he following approximate values a r e quite sa t is fac tory

Equation (5) becomes

log K265 = 2 log mK+ + 2 log y K + + 2 pH (7)

Digression No. 3

Equation (5) contains a pH t e r m (pH = -log aH+). Because the liquid sample taken a t t h e su r face is greatly depleted in C 0 2 and H2S re la t ive t o the reservoir fluid, t h e pH measured on t h e laboratory sample is really of no use t o us and we need to ca lcula te a new value, pHt, for t h e reservoir fluid a t 265 '~ . This is not a trivlal task and i t requires an i tera t ive calculation best performed on a computer or programable calculator. The usual procedure involves using an ion or proton balance fo r a l l t he pH-sensitive ions in solution. For more details, s e e Henley e t al. (1984). Through these methods, pH265 for the BR22 discharge is found t o be 6.1.

A simplistic calculation o f t en provides a useful approximation of pH. Consider t h e equilibrium

where H ~ C O ~ * represents undissociated dissolved co2.

A t 2 5 ' ~ log K = - 6.36 so t h a t in t h e liquid analysis we may calcula te (with yHCO3- '0.8 a t laboratory temperature) t h a t

3 mHco3- = 3.29 x 10- and mH2C03* = 0.25;

Recalculating t h e discharge composition thro gh (21, a t 26i°C, mHco3- = 2.67 x 10-Y and

mH2C03* = 101.2 x Log K fo r reaction (8)

is -7.84 a t 2 6 5 ' ~ and s ince

log K = log m H c 0 3 - + log YHco3- + log a ~ +

- log mH2C03+ we may subst i tu te values t o obta in pH. From these d a t a our pH es t ima te is 6.1! (There a r e a number of hidden assumptions in this approach so t h a t i t may

TEMPERATURE O C

Figure 2.4. Experimental solubility data for quartz in water as a function of temperature.

be qui te er roneous for s o m e wells, particularly those discharging f rom a heavily e x p l o ~ t e d reservoir.)

Giggenbach (1980) has developed an approach t o mineral-fluid equilibria which avoids t h e need t o conserve A1 a s required in reaction (4) or

+'Ye p+H itself. In th i s approach t h e log ra t io A1 / a designated AIH, is used a s a principal variable. #i: makes fo r fun reading a s you can imagine - but t h e method is ext remely valuable!

The log analyt ica l quot ient a K + / a ,. . ,Or the reservoir fluid is -2.56 + 6.1 = 3.54. his value, compared t o t h e equilibrium quotient (3.94), suggests t h a t t he reservoir fluid l ies in t h e 'mica' stabil i ty field, 0.4 log units away f rom t h e mica-Kspar boundary. This is compatible with the petrography since sericit ization is widespread, but a r e the re some d a t a checks t o be made?

1. Experimental determinat ion of f r e e energies for t h e r eac tan t s and products, particularly ions, i s very difficult . Direct determination of the equilibrium constant is equally difficult due t o unfavorable kinetics. The reaction coefficients required in hydrolysis reactions also require ex t r eme precision in t h e thermodynamic data.

2. A t 2 6 5 O ~ , ser ic i te is be t t e r described a s i l l i te and the ac t iv i ty of the KA13Si301, (OH)2 component should be determined, if possfile, by x-ray and this value added into equation 7.

3. The pHt calculation is dependent upon precise thermodynamic d a t a for ion association reactions and reliable e s t ima tes of ac t iv i ty coefficients.

L O G

aNo;/aH+

Exercise 4. Is t he reservoir fluid close to equilibrium with respect t o Kmica and Kspar a t qua r t z sa tura t ion o r is t he pHt too high or low?

-

-

'- -

2-

-

Other hydrolysis react ions may be assessed similarly and t h e results a r e usually expressed on act iv i ty r a t io diagrams such a s those shown in Figures 2.5 and 2.6.

Exercise 5. Loca te the BR22 fluid composition point in Figures 2.5 and 2.6 using t h e d a t a of Table 2.3 and the e s t ima ted reservoir fluid pHt.

I I 1

+QUARTZ 2 5 0 " C

ALBITE

The calcium aluminosil icates a r e a particular problem because they occur in solid solution with iron and because thei r thermodynamic da ta a r e poorly known. For th is reason Figure 2.5 has been const ructed using field da ta for t he calcium aluminosil icate phase boundaries.

(L a a. V) a -I W . L

I x

MONT.

Exercise 6. Scan the introductory section and decide which is t h e dominant ca lc ic phase in t h e a l tera t ion assemblage.

KAOLlNlTE

Digression No. 4.

The mineral is, of course, ca lc i te , which, due t o the high C 0 2 con ten t of t h e fluid, f ixes most of t h e available Ca". To assess i t s solubility in t e rms of aCa++/a2H+ you will need an appropriate log K d e r ~ v e d f rom Appendix Table 2.A1.

K-MICA

L O G a,,+ aR+ /

2 4 6

LOG a K/ + a,,+

F i g u r e 2.5. A c t i v i t y - a c t i v i t y r e l a t i o n s f o r the F i g u r e 2.6. A c t i v i t y - a c t i v i t y r e l a t i o n s f o r the s y s t e m Na20-K20-A1203Si02-H20. s y s t e m Ca0-K20-A1203Si02-H20.

R. W. HENLEY & K. L. BROWN 3 1

- Te K o p l a ( O . O O l )

C a C 0 3 + 2H+ = ~ a + + + H2c03* 1

The coeff ic ients of t h e regression equations (Appendix Table 2.AI) may be summed t o give 7-

I loe K = P

\ " -235.56 + 10676.81/T + 89.084 log T - 0.0392 T g .- e K a w e r a u ( O . 0 3 ) -

At 2 6 5 ' ~ th is equation gives log K = 6 46. The 5 analytical quotient, ac ++mH2C03*/a2H', for t he m 6. BR22 fluid is 6.1, whit% suggests t h a t t h e fluid is ~ ~ h i ~ ( 0 . 1 )

close t o sa tura t ion with calcite. The di f ference between t h e analytical and equilibrium quotients is \

C e r r o P r l e t o ( 0 . 2 5 ) 0.36 log units, a small value compared with the sum of possible e r ro r s on p ~ , ~ C a + + , and, if underground boiling has occurred, mH2C03*.

1

D i g r e s s i o n No. 5. 200 300

Temperature "C

Control of fluid pH through sil icate react ions has an Figure 2.7. Relationship between fluid pH and important implication. Taking t h e Kspar-Kmica salinity for fluids in equilibrium with reaction, fo r example Kspar-Kmica-quartz.

i t is self-evident t h a t mK+ is a function of pH. Potassium and sodium a r e the predominant cations in these high-temperature solutions and thei r r a t io is l imited by t h e react ion of plagioclase to Kspar and Kmica, and, below about 2 7 0 ' ~ ~ the format ion of secondary albite. The l a t t e r , t h e coprecipitation of a lb i te and Kspar, af fords a l imiting condition for t h e ra t io of sodium and potassium, which is used a s a geothermometer. Combining the two reaction constants, a t empera tu re dependent function of pH and m ( ~ K) is found, and since (mNa + m ) mcl-, PA is re la ted t o salinity. Flgure &' shows thls function together with some da ta points for o ther ac t ive geothermal systems.

From a n assessment of fluid compositions from a number of wells, Giggenbach (1981) showed t h a t t h e Kspar-Kmica react ion was important in the upper two-phase (i.e., boiling ) zone of geothermal systems, but t h a t in general another equilibrium prevailed

plagioclase + H 2 c 0 3 * = 'clay' + calc i te (9)

Da ta from a number of New Zealand fields a r e summarized in this context in Figure 2.8. Giggenbach (1980) has also shown t h a t i t may be more appropr ia te t o r e fe r t o geothermal mineral- fluid in teract ions with reference t o a s teady s t a t e ra ther than t o equilibrium, but this does not a f f e c t t h e validity of t h e general relations discussed here.

Another important group of consti tuents ( the gases ammonia, methane, nitrogen, hydrogen, etc.) a r e included in Tables 2.2 and 2.3. These d a t a provide a g rea t deal of information about sources of various components, processes in t h e aquifer, etc.--see, for example, Giggenbach (1977, 1980). Although in the laboratory gas + gas react ions a r e normally slow, they appear t o be a s f a s t a s silica precipitation in t h e fluid + mineral sys tem of a geothermal aquifer. This provides a useful independent check on aquifer t empera tu res through reactions l ike

although if t h e fluid boils e n rou te t o t h e well vapor loss o r gain may occur in the discharge re la t ive t o t h e aquifer fluid. For th is well, BR22, these e f f ec t s a r e minor and t h e analysis quotients for these react ions a r e close t o those expected for equilibrium a t 265'C.

Reviewing d a t a from Broadlands and other fields, Giggenbach (1980) showed t h a t t h e ra t io H2/H2S was a function of t empera tu re (Fig. 2.9) ref lec t ing equilibrium between pyrite and Fe-silicates represented by chlor i te and/or epidote. Only near 300°C do t h e observed hydrogen contents approach those for t h e equilibrium between pyrite and pyrrhotite, or pyrrhotite + Fe-silicate.

o Woirokei o Kawerau

Broodlands o other oreas

Figure 2.8. Boiling point vs depth (a) and fluid pressure P (b) curves for a series of mineral assemblages. The data points for Wairakei, Kawerau, Broadlands, El Tatio, and Iceland wells correspond to the point at which measured temperature/depth and temperature/pressure curves changed from (nearly) constant temperature with depth behaviour to that indicating boiling point vs depth conditions (reproduced with permission from Giggenbach, 1981).

FLUID-MINERAL EQUILIBRIA: TRACE-METAL CONTENTS

Table 2.4a summarizes base- and precious-metal analysis from waters discharged a t a tmospher ic pressure from two Broadlands wells (Weissberg e t al., 1979).

As noted earlier, base-metal sulfides a r e t o be found in drill co re and cuttings from many of t h e wells a t Broadlands, but gold has been repor ted only in a complex antimony-arsenic-mercury-thallium-sulfide precipi ta te in a hot spring. F r e e gold (as e l ec t rum)

Figure 2.9. The ratio H2/H2s as a function of temperature for selected mineral pairs in relation to data from active geothermal systems, plotted with respect to quartz and alkali geothermometer temperatures. (FeO) represents an iron-aluminium silicate (chlorite?, epidote?). CH4/Co2 contours are also shown. Note the differences between silica and alkali geothermometer temperatures (T is the NaKCa empirical geothermometer oYK$ournier and Truesdell, 1973). These differences reflect aquifer processes consequent on exploration and have high significance for monitoring geothermal field development (reproduced with permission from Giggenbach, 1980).

does occur in co re f rom nearby systems, e.g, Kawerau (B. Christenson, personal communication) and, s ee below, is t o be ant ic ipated in Broadlands core--but no one has yet looked for it.

Recen t work a t BR22 has shown t h a t the concentrations of t r ace meta ls discharged a t a tmospher ic conditions do not always represent t he concentrations of those me ta l s in t h e deep geothermal fluid (Brown, 1985). Considerable deposition occurs in t h e piping so t h a t corrected concentra t ions (Table 2.4b) a r e in f a c t much higher than would be judged f rom Table 2.4a.

Table 2.4a--Concentration of t r a c e m e t a l s i n w a t e r d i s c h a r g e d a t t h e s u r f a c e from e x p l o r a t i o n w e l l BR2 a t Broadlands

Table 2.4b--Minimum c o n c e n t r a t i o n s of t r a c e and b a s e m e t a l s i n w a t e r d i s c h a r g e d from BR22 a t Broadlands ( ~ r o w n , 1985). These d a t a a r e based on e m p i r i c a l d a t a f o r t h e m e t a l c o n t e n t s of t h e t o t a l d i s c h a r g e , b u t f o r e a s e of compar ison w i t h Table 2.4a, t h e s e d a t a a r e c o r r e c t e d t o 1 b.a. s e p a r a t i o n p r e s s u r e , a l t h o u g h a s d i s c u s s e d below, a l a r g e p e r c e n t a g e of t h e coppe r , s i l v e r and gold i s p r e c i p i t a t e d p r i o r t o d i s c h a r g e a t a tmosphe r i c p r e s s u r e .

S e p a r a t i o n P r e s s u r e Fe Cu P b Zn Ag Au

Exercise 7. Calcula te t h e aquifer concentra t ions of these me ta l s a t BR22. What percentage of t he gold and silver i s prec ip i ta ted by boiling during discharge t o t h e surface? Water from BR22 flashed t o 1 bar a. contains similar Au and Ag t o those shown in Table 2.4a fo r BR2.

The concentra t ion of. these me ta l s may be contro l led by mineral deposition o r dissolution in t h e aquifer , so f i r s t we must examine, a s with ca l c i t e and t h e various si l icates, whether t h e fluids a r e c lose t o sa tura t ion with respect t o common sulfides, or t h e nat ive me ta l (for gold).

Lead - Figure 2.10 shows the distribution of t h e

d i f f e r en t lead chloride species a s a funct ion of m based on r ecen t exper imenta l da t a (Seward, 19&( F rom these d a t a we s e e t h a t PbCl const i tu tes about 50% of t h e lead in solution a s a c t l o r ide complex a t 2 6 5 ' ~ in a low salinity fluid. The solubility of galena may then be wr i t t en

PbS + 2Hf + X I - = PbC12 + H2Saq ( 1 0 )

Insert ing analytical d a t a for t h e reservoir f luid yields m p ,i 0.015 ug/kg (as PbC12), so t h a t t h e a c t u a l so u b l l ~ t y of galena is, including a l l t he chloride species, 0.03 ug/kg. This value is much lower t han t h e 1-10 kg actual ly found in these wa te r s and indica tes

t h a t in t hese low sa l in i ty fluids, some o the r complex species mus t be present; however, t h e r e a r e a s ye t no exper imenta l d a t a available fo r possible species like P ~ ( o H ) + , ~ b H C 0 3 , etc., t o ca l cu l a t e which predominates under t hese conditions.

Digression No. 6

Figure 2.11 shows t h e ca lcula ted solubil i ty of galena a s a function of t empera tu re fo r a 1.0 m chloride solution (Q6 wt% NaCI). What is t h e increase in solubility t o be expec ted a t 2 6 5 ' ~ if species like PbHCO; a r e present? Fi rs t wr i t e a s imple react ion fo r t h e format ion of P~HCO;.

PbS + 2H+ + HCO? = P~HCO; + H2S (11)

Increasing chloride re la t ive t o t h e Broadlands fluid does not a f f e c t t h e react ion a s wr i t t en excep t through changes in ac t iv i ty coefficients, so t h a t t h e observed contribution due t o th is o r o the r species may be p lo t ted d i rec t ly on to Figure 2.10. I t is c l ea r t h a t t h e chloride complexes dominate a t t he se higher salinities. Not ice a lso t h a t in a hydrothermal sys tem with fluid a t th is salinity, t h e pHt of t h e fluid would be about 1 t o 1.5 units less than fo r Broadlands (Fig. 2.7). The increased solubility due t o th is e f f ec t , reac t ion (101, is, however, part ly of fse t by t h e decrease of m H C O j

Incidentally, using d a t a f rom Table 2.A1 and t h e observed lead con ten t of t h e Broadlands fluid, a n e s t i m a t e of log K265 fo r th is reac t ion is 7.58.

Gold

In con t r a s t t o lead (and zinc), me ta l s such a s gold, a rsenic and ant imony have very s table bisulfide complexes. Fo r gold t h e principle dissolution react ion (Seward, 1973) is

and a regression equation fo r t h e equilibrium constants of this reac t ion i s given in Appendix Table 2.A1.

Digression No. 7

Determinat ion of t he solubil i ty of gold introduces a new solution pa rame te r , t h e re la t ive redox s t a t e , which is tradit ionally represented by t h e value of log fO2 fo r t h e solution, where f i s t h e fugacity. The choice of fO2 is unfor tunate since i t s fugaci ty is immeasurably small , whereas fH2 can be more easily directly measured. Since gas concentra t ions a r e relatively low, i t is permissible t o subs t i tu te t h e partial pressure, Pi, f o r fugacity.

The concentra t ion of H 'n t h e reservoir fluid f rom Table 2.3 is 1.07 x 1 6 ' molesikg. In order t o ca lcula te fy2,. a knowledge of Henry's Law is required. Thls IS given by

F i g u r e 2.10. T h e p e r c e n t a g e d i s t r i b u t i o n o f lead as c h l o r o l e a d ( I 1 ) complexes as a f u n c t i o n o f to ta l chloride c o n c e n t r a t i o n , CIT, u p t o 3 0 0 ~ ~ a t the s a t u r a t e d v a p o r p r e s s u r e o f the s y s t e m ; the c u r v e f o r each c o m p l e x is la- belled a c c o r d i n g t o l i g - a n d n u m b e r , e.g., c u r v e s l a b e l l e d ' 2 ' r e f e r t o p b c l Z O ( r e p r o d u c e d w i t h permission f r o m S e w a r d , 1984) .

P i = X i s where s ti = Henry ' s C o n s t a n t , 9 f o r i

Xi = mole f r a c t i o n of i

P i = P a r t i a l P r e s s u r e of i

The mole f r a c t i o n o f hydrogen i n t h e r e s e r v o i r f l u i d i s

Henry's cons tant fo r H2 a t 2 6 5 ' ~ is 26,000 bars/mole-fraction, and the re fo re

f H 2 = 0.182 X X 26000 = 0.0473

A t 2 6 5 ' ~ log K = -19.46 f o r t h e react ion

H20(1) = + 1/202(g)

Therefore , log PO2 = -36.27.

L o c a t e t h e BR22 reservoir fluid on t h e f02/pH diagram provided in Figure 2.12.

A t t h e t empera tu re s under discussion for Broadlands (<300°C), magne t i t e is no t present a s an a l tera t ion mineral. I t s s tabi l i ty f ield s eems t o be usurped by

- mc1- = 1 - ) rnCl-=O.5 - - Soluhili t y decrease

m 0 - Y

2 - 2 .

D - 4 - a m 2 - 6 ,

1 / B = A x conduct~ve coolinq C = A x adiabatic cooling

Figure 2.11. Solubility of galena in aqueous chloride solution.

t h a t of chlorite, a f ea tu re noted in o ther sys tems and in many hydrothermal deposits (Barton e t al., 1977).

Figure 2.12 contains solubility contours for AU(HS)~-.

Exercise 8. Compare the measured gold content of t he Broadlands fluid with the solubility contours given in Figure 2.12. What conclusion may b e drawn? Discount any suggestion of chloride complexes which a r e precluded (a t this temperature , salinity, and fO2) on t h e b a s i s of exper imenta l da t a (Henley, 1984, unpublished data). Sketch isosolubility curves (lead, mg/kg) in t h e H2S field for galena a s chloride complexes alone and no te how differently they behave compared t o those of gold.

Other Metals: Copper, Silver, and Arsenic

Other me ta l s may be t r e a t e d in t h e s a m e manner although i t may be necessary t o consider a wider range of complexes. Available d a t a for copper give t h e following species distribution in t h e Broadlands aquifer fluid a t 2 6 5 ' ~ (Table 2.51, although the re is some doubt a s t o t h e precision of t h e format ion constants for t he chloride species.

By f a r t h e predominant species i s t h e Cu(HS); complex and the concentration of th is species i s essentially equal t o the measured aquifer concentration in Table 2.4b.

using t h e d a t a below and tha t of Table 2.5.

What a r e t h e predominant copper species here?

For silver, only d a t a for t h e chloride complexes have Seen measured (Seward, 1976).

Ag2S(S) + 2 ~ ' + 2C1-

= 2AgCI + HZS(aq); log K2650 = -5.5 (13)

Ag2S(s) + 2 ~ ' + 4C1-

= 2AgC12 + H2S(aq); log K2650 = -2.1 (14)

From these data , t h e solubility of a rgen t i t e is

~ o - ~ . ' O + ( 1 0 - ~ . O ~ / . 6 9 4 )

= 1.965 x m l e s / k g = 0 . 2 1 p p b .

I TOTAL S 2 x 1 0 - 3 m

I t

Exercise 9. By assuming coexistence of t h e Same Figure 2-12, ~ o g fo2 versus p~ diagram for sulfides (Table 2.51, e s t ima te t h e fluids and alteration minerals at Broad- concentra t ions of CuC12 and Cu(HS)- lands at 2 6 0 ~ ~ . Superimposed are solubili- fo r t h e Salton Sea geothermal f lu i3 ty contours for gold in mg/kg.

Table 2.5--Distribution of copper bisulfide and chloride complexes in the BR22 aquifer fluid at 265'~. Thermodynamic data from Crerar and Barnes (1976) for the b o r n i t e - c h a l c o p y r i t e - p y r i t e coexistence boundary. Units; g Cu/kg.

Exercise 10. Compare th i s value with t h e measured value in Table 2.4b. What conclusions may b e drawn?

Assuming t h a t si lver forms a n Ag(HS)? complex t o accoun t f o r t h e e x t r a solubility, ca lcula te a value fo r log K f o r t h e solubility of Ag2S in t h e aquifer.

In a l l t h e New Zealand geothermal sys t ems arsenic concentra t ions a r e high r e l a t i ve t o o the r na tu ra l waters. The hydrothermal chemis t ry of arsenic is poorly known. Weissberg e t al. (1966) found high solubil i t ies fo r orpiment ( A S ~ S ~ ) in alkaline solutions up t o a o u t 200°C. Arseni te ( ~ s 0 3 ~ - ) o r th ioarseni te (ASS>-) complexes a r e t h e most likely t o account f o r arsenlc t ranspor t in t hese fluids. Similarly t h e chemis t ry of ant imony is poorly known. Elementa l mercury has a re la t ive ly high volati l i ty and may be significantly t ranspor ted in any vapor f rac t ion fo rmed by boiling, however i t s speciation in solution is a lso very poorly understood a t high tempera tures .

MINERAL DEPOSITION

Silica

Whereas in t h e reservoir we showed t h a t t h e si l ica concentra t ion of t h e fluid was controlled by t h e dissolution of quar tz , a t lower t empera tu re s ( 150°) si l ica solubility is controlled by t h e equilibrium with amorphous silica. The solubilit of amorphous si l ica in 8 . . t h e t empera tu re range 0 t o 250 C IS glven by (Fournier and Rowe, 1977).

log S = -731/T + 4.52

where S = solubility in mg/kg, T = absolute t empera tu re (OK).

When geothermal water ar r ives a t t h e su r f ace through a well or fissure, i t becomes supersa tura ted with r e spec t t o amorphous si l ica due t o both t h e reduction in tempera ture , and t h e concentra t in e f f e c t % of s t e a m loss. A t BR22 fo r example, a t 98 C, t h e si l ica concentra t ion in t h e separa ted wa te r is about 800 mg/kg. The solubility of amorphous si l ica i s only 355 mg/kg allowing possible deposition of a lmost 0.5 kg si l ica per t o n of fluid. However, only a very sma l l pe rcen tage of t h e excess silica is deposited. Most of t h e exces s si l ica polymerizes t o form a colloidal suspension (P la t e 2.IA)--a process t h a t may t a k e severa l hours. Subsequent growth and aggregat ion of

t h e colloidal par t ic les fo rms si l ica s in ter (P l a t e s 2.18 and 2.IC). If t h e par t ic les a r e ordered and a l l t h e s a m e size, then gem quali ty opal is formed. Deposition processes fo r si l ica a r e considered in much more deta i l by Fournier (1985, th is volume).

Ca lc i t e

Ca lc i t e somet imes fo rms inside t h e cas ing of geothermal wells during d ischarge and, a s we discovered ear l ie r , t h e reservoi r fluid is s a t u r a t e d with respect t o calcite. The equil ibrium re la t ion may be wri t ten

Exercise 11. Derive a regression equat ion fo r t h e t empera tu re dependence of reac t ion (151, using d a t a f rom Appendix Table 2.AI. Your resul t should be

log K = -226.43 + 6552.81/T + 89.084 log T - 0.0746 T

Since each mole of ca l c i t e t h a t dissolves in react ion (15) produces one mole of Ca++ a n d t w o of H C O j , t h e equilibrium cons t an t can be modified t o read

With these two expressions we can draw solubility curves f o r ca l c i t e a s a function of P and t. Such a s e t of curves is shown in ~ i ~ u $ ~ $ . l 3 based on exper imenta l data. Notice t h a t ca l c i t e is more soluble a t low t empera tu re t han high tempera ture .

Digression 8

The to t a l fluid pressure in a n aquifer is t h e sum of t h e water and gas pressures. PCO2 may be ca lcula ted f rom Henry's Law

H2C03+ = C 0 2 , g + H 2 0 (17)

A t 2 6 5 ' ~ ~ log K = 1.917 = log PCO2/mCO 2

log PCo2 = 1.917 + log (0.104)

PC02 = 8.5 bars, and with P H = 50.8 bars, 2

Preservoir = 59.3 bars

R. W. HENLEY & K. L. BROWN 3 7

F i g u r e 2.13. E x p e r i m e n t a l l y d e t e r m i n e d s o l u b i l - i t y o f calcite; n o t i c e the e f f e c t o f d is - s o l v e d sa l t s ( r e p r o d u c e d w i t h p e r m i s s i o n f rom E l l i s , 1963).

This, incidentally, means t h a t t he boiling-point depth for th is solution is about 100 m deeper than fo r pure water, but that's a whole s to ry in itself (Hedenquist and Henley, 1985). Figure 2.8b shows t h e e f f e c t of C 0 2 on boiling depth where PC02 is buffered by the plagioclase, clay, ca l c i t e r e a c t ~ o n . Consider a well discharging from this aquifer. If t h e wellhead pressure is held a t 15 bars, t h e fluid experiences a pressure drop of nearly 45 bars during flow up t h e well. Since C 0 2 is a relatively volati le species, i t parti t ions strongly in to any coexisting vapor phase a s is clearly demonstra ted by t h e original analytical data. The concentra t ion of C 0 2 in t h e liquid phase a t equilibrium with vapor may be obtained from the following equation

where B is t h e parti t ion coeff ic ient for C 0 2 between vapor and liquid. Regression equations for B a r e given in Appendix Table 2.A1.

Exercise 12. From (16) and (171, ca lcula te the mH2Co3+ and PCO2 aq of t h e liquid phase resulting from 'steam format ion a t 2 2 0 ' ~ a s t h e well discharges, and then t h e s t a t e of sa tura t ion of ca l c i t e a t this temperature .

You should find t h a t (a) a t 220°C t h e aqueous phase is about 4 t imes supersa tura ted with ca l c i t e and (b) each ton of fluid passing up t h e well could deposit 30 gm of ca lc i te ( 15 cc). What would be t h e supersaturation a t 2oo0C?

Digression No. 9.

The pH of the liquid sepa ra t ed a t t h e wellhead a t 184OC is 7.3 (this is a value recalcula ted f rom t h e analytical da t a a s briefly outlined in Digression 3). Why is t he flashed fluid more alkaline than t h e reservoir fluid?

Consider t h e reaction

Clearly, even without considering t h e change of K for this react ion due t o t empera tu re drop, removal of C 0 2 by boiling mus t result in a pH increase. ~emova 'Fof 90% of t h e C 0 2 increases pHt by 1 unit.

Metal Sulfides and Gold

Galena, sphalerite and chalcopyr i te somet imes occur within the ca l c i t e scale coat ing t h e insides of well casings. Metal sulfides have also been precipi ta ted in t h e surface installations. Figure 2.14 is an x-ray diffraction pa t t e rn for a s ca l e deposited inside wellhead equipment a t BR27 and similar scales occur a t aR22. Galena, pyrite, chalcopyrite, and sphalerite a r e all present. There a r e also peaks due t o t h e presence of e lec t rum. A scanning e lec t ron micrograph of t h e scale i s shown in P l a t e 2.2. The scale was deposited in the wellhead plumbing where t h e pressure drops from -40 bars t o 11 bars with a concomitant t empera tu re drop t o 185 '~ . The boiling t h a t accompanies t h e drop in pressure has a number of ef fects . As shown above, C 0 2 parti t ions strongly in to t h e vapor phase, so t h a t t h e llquld phase experiences a rise in pH. H2S is also lost f rom t h e solution, so t h a t reactions such as:

a r e driven t o the l e f t leading t o gold deposition. The equilibrium constant for this react ion a t 1 8 5 ' ~ is log K = -8.4 (Appendix Table 2.A1). Using t h e distribution coefficient, B = 174, f rom Appendix Table 2.A1 and the measured H2S concentration of 0.103 moles/kg of s team, we can calcula te

Before we can calcula te t h e solubility of gold from reaction (201, we need a value fo r PHZ f rom the separa ted liquid. There a r e a number of problems involved in es t imat ing PH during irreversible processes such a s flash boiling t ~ h o r s t e n s e n , 19841, but one guess may be obtained f rom t h e reaction

Br 27

ORIFICE PLATE

Figure 2.14. X-ray diffraction pattern of material deposited at BR22 well. C = chalcopyrite, G = galena, P = pyrite, M = magnetite, E = electrum.

Substi tuting values fo r ~ H S ? HS, mSpt, ,SO;, and pH, 1.9 &/kg a t 185 '~ . When this value is subs t i tu ted in assuming t h a t t h e react ion 1s a t equl ~ b r l u m , t h a t reac t ion (20) together with t h e H2S act iv i ty , log P H no oxidation of H2S t o SO; had occurred in t h e sample , -2.85. This is in remarkable ag reemen t with t he vafu; then using values f rom Appendix Table 2.AI. of -2.72 derived f rom react ion (21) indicating t h a t t h e

"true" or e f f ec t ive log P H Z must l ie around these log PH2 = -2.72 values a t leas t a s f a r a s gold 1s concerned.

An a l t e rna t ive method for es t imat ing PH2 is by using t h e measured par t ia l pressure of hydrogen. The amount of hydrogen sepa ra t ed a t 11 bars i n to t h e gas phase is 0.57 x moles/kg (from Table 2.2). The partial pressure of hydrogen (PH2) is given by the product of t h e mole-fraction of H2 and t h e to ta l pressure

Exercise 13. Compare t h e e f f e c t s of boiling on dissolved me ta l con ten t for lead, silver and copper.

The discussion of arsenic and antimony sulfide deposit ion is hindered by t h e absence of chemical d a t a fo r t h e complexing of t hese metals. Natura l occurrences (Weissberg, 1969; Hedenquist and Henley, 1985) and deposits formed during chemical processing of geothermal discharges suggest t h a t pH is a controll ing var iable through a react ion of t he form

This value of pH2 places t h e solution outside t h e stabil i ty f ield for H2S which clearly from t h e analysis is s table in t h e solutlon col lec ted a t I I bars. I t would, - therefore , appea r t h a t th is and o the r compet ing redox react ions in t h e solution have not c o m e t o equilibrium in t h e new pressure regime.

The measured gold concentra t ion of 2.3 pg/kg (Table 2.4b) a t I bar corresponds t o a concentra t ion of

Observations on arsenic deposit ing in hot springs and well discharges (Hedenquist and Henley, 1985) suggest t h a t local deposition may occur where normal near neut ra l pH fluids cool without loss of C 0 2 so t h a t pH is buffered by t h e bicarbonate-carbonic ac ld pair o r where acidification occurs, perhaps by mixing with surficial ac id wa te r s a s a t Tamagawa Springs, J apan (Nakagawa, 1971).

Plate 2.1. a). Scanning electrron micrograph of colloidal silica from BR22 w e l l after initial polymerization. b). After further growth an6 agglomeration - note the porous structure - and c). the depsited scale with the pres filled wit3 silica to form a hard vitrmus sinter.

Plate 2.2. Scaaning electron xicrograph of Broadland's sulfide scale. The x-ray diffraction pattern for this naterial is shown in Figure 2.14. Each segment af the bar corresponds to 50 and, for convenience, conversion to local T z ~ p w.i.ts is provided.

F i g u r e 2.15. D i s t r i b u t i o n o f z i n c chloride s p e c i e s a s a f u n c t i o n o f chloride c o n c e n t r a t i o n a n d t e m p e r a t u r e i n a c i d i c s o l u - t i o n s . T h e n u m b e r s s h o w n are the number o f c h l o r i d e l i g a n d s f o r each z i n c a tom ( r ep roduced w i t h p e r m i s s i o n f rom Ruaya and Seward, in p r e s s ) .

ACKNOWLEDGMENTS

We a r e gra teful t o Bruce Christenson fo r reviewing t h e manuscript , and t o Sharon Thorne fo r cheerfully and painstakingly preparing t h e manuscript . T. M. Seward and Ruaya kindly allowed prepublication reproduction of t h e Zn complexing data. One of us (KLB) found no solace a t a l l in t h e accompaniment of J a m e s Galway.

REVIEW QUESTIONS

I. Well BR22 discharges about 100 tonnesihr; how long would i t t a k e t o deposit a million ounces (2.5 x lo6 gm) of gold a t t h e wellhead? What does your answer imply about t h e chemis t ry of t h e fluids responsible fo r epi thermal gold deposits? If a l l t h e potential silica and ca l c i t e coprecip i ta ted with gold - and we continuously c leared t h e well! - what would be t h e gold concentra t ion of t h e deposit?

2. From Figure 2.2 e s t i m a t e t h e mass f rac t ion of vapor formed p e r lo°C t empera tu re drop a s liquid (initially on t h e two-phase boundar ) adiabatically d boils f rom 3 0 0 ' ~ and 250°C t o 200 C.

3. What d i f ferences do you recognize between t h e processes within a discharging well and in a fluid moving upward through a fissure t o t h e surface?

4. F igure 2.15 shows t h e distribution of z inc chloride complexes a s a function of chloride con ten t and tempera ture . The equilibrium constant (as log K) fo r t h e dissolution of ZnS

ZnS + ZH* = Zn2+ + H2Saq

equals -2.7 (Helgeson, 1969) a t 2 5 0 ' ~ and -2.23 a t 300°C. Ca lcu la t e t h e solubility of sphaler i te in t h e fluid discharged f rom Broadlands, BR22, and compare i t with t h e observed z inc con ten t of t h e

log C, ( m o l l k g )

fluid. What conclusion may be drawn about z inc t ranspor t in t h e Broadlands aquifer fluid? Make a similar calculation for t h e Salton Sea brine using d a t a provided earlier .

5. In Kuroko deposits (submarine polymetall ic massive sulfides), m e t a l sulfides a r e thought t o be prec ip i ta ted a s t h e submarine hydrothermal f luid ( 3 0 0 ~ ~ ) mixes with cold seawater . Consider t h e chemis t ry of o re deposition a s a resul t of cooling conductively or mixing with seawater . Are t h e r e any other possible depositional mechanisms?

6. Apar t f rom boiling, what o ther prec ip i ta t ion mechanisms can you propose f o r silver and gold in epi thermal deposits?

7. Consider t h e e f f e c t of salinity on t h e solubil i t ies and deposit ional mechanisms of gold, silver, l e ad and z inc in hydrothermal fluids. Is dilution by cold, f resh wa te r a process which could prec ip i ta te any of t hese meta ls?

8. In a n epi thermal vein adular ia and ca l c i t e a r e found. Would th is paragenesis provide informat ion on physical conditions during o r e deposition?

REFERENCES

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Brown, K. L., 1985, Gold deposition f rom New Zealand geothermal wells: Economic Geology (in press).

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a t 200° t o 350°C: Economic Geology, v. 71, p. 772-794.

Ellis, A. J., 1963, Solubil i ty of ca l c i t e in sodium chloride solutions at high tempera tures : American Journal of Science , v. 261, p. 259-267.

Ewers, G. R., and Keays, R. R., 1977, Volatile and precious-metal zoning in t h e Broadlands geo the rma l field, New Zealand: Economic Geology, v. 72, p. 1337-1354.

Fisher, R. G., 1959, T h e na tu ra l h e a t flow f rom t h e upper Waiora Valley: Geophysics Division, DSIR, Geo the rma l Circular TSG 5.

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Fournier, R. O., and Po t t e r , R. W., 1982, A revised and expanded si l ica (quartz) geothermometer : Geothermal Resources Council Bulletin, v. I!, p. 3-9.

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Helgeson, H. C., 1969, Thermodynamics of hydrothermal sys t ems a t e l eva t ed t empera tu re s and pressures: American Journal of Science, v. 267, p. 729-804.

Henley, R. W., Truesdell , A. H., and Barton, P. B., Jr., 1984, Fluid-Mineral Equilibria in Hydrothermal Systems: Socie ty of Economic Geologists, Reviews in Economic Geology, v. I , , 267 p.

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Morey, G. W., Fournier, R. O., and Rowe, J . J., 1962, The solubility of q u a r t z in wa te r in t h e t empera tu re in terval f rom 25 t o 300'~: Geochimica e t Cosrnochimica Ac ta , v. 26, p. 1029-1043.

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Seward, T. M., 1973, Thiocomplexes of gold and t h e t ranspor t of gold in hydrothermal o r e solutions: Geochimica e t Cosmochimica Ac ta , v. 48, p. 121-134.

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Appendix Tab le 2.Al--Summary o f thermodynamic d a t a

React i o n Equa t ion Data

Temp. Range Refe rence

1. S i 0 2 , + 2H20 = Si(OHl4 l o g K = 0.104 - 1 1 6 2 . 8 7 1 ~ Y q

2. 3KA1Si308 + 2Hf = K A 1 3 ~ i 3 0 1 0 ( ~ ~ ) 2 log K = 5.1062 + 6.31 x

+ 6S i02 ,qz + 2 ~ ' ~ o - ~ T + 1302.51T

3. H2co3* = HCOj + H+ l og K = 6.2 -1.897 x T

-2.062 x 103/T

4. CaC03 + H2co3* = ca++ + 2 ~ ~ 0 ; l og K = -223.16 + 6552.811T

+ 89.084 l o g T - 0.0771 T

5. H2co3* + H20 l o g KH = 3.2702 - 0.002515 T

6. PbS + 2 ~ 1 - + 2 ~ + = PbC12 + H2Saq l o g K = 19.417 - 9 0 3 0 . 1 / ~

7. Au + 2H2S = A U ( H S ) ~ - + H+ + 112 H2 l og K = 834.443 - 3 6 9 4 5 . 1 7 1 ~

-305.449 l o g T + 0.11046 T

8. H2Saq = HS- + H+ log K = -12.18 + 2377.51T

9. H201 = H2,g + 1 /202 ,g l o g K = 7.6 - 1 4 5 6 4 . 1 3 1 ~

10. HS- + 202 = so4= + H+ l o g K = -31.75 + 49975.51T

P a r t i t i o n c o e f f i c i e n t s ( f o r p u r e w a t e r s o l v e n t )

11. C 0 2 ,,/C02, 1 l og B = 4.7593 - 0.01092 t

12. H2S,,/H2S, 1 l o g B = 4.0547 - 0.00981 t

Refe rence l ist f o r Appendix Table 2.Al

(1) Fournier, R. 0.) and Po t t e r , R. W., 1982, A revised and expanded silica (quartz) geothermometer : Geothermal Resources Council Bulletin, v. I I , p. 3-9.

(2) Helgeson, H. C., 1969, Thermodynamics of hydrothermal sys tems a t e levated t empera tu re s and pressures: American Journal of Science, v. 267, p. 729-804.

( 5 ) Ellis, A. J., and Golding, R. M., 1963, The solubility of carbon dioxide above IOOOC in water and in sodium chloride solutions: American Journal of Science, V. 261, p. 47-60.

( 6 ) Seward, T. M., 1984, The format ion of lead (11) chloride complexes t o 300'~: Geochimica e t Cosmochimica Acta, v. 48, p. 121-134.

(3) Read, A. J., 1975, The f i rs t ionization constant (7) Seward, T. M., 1973, Thiocomplexes of gold and of carbonic ac id f rom 25 t o 2 5 0 ' ~ and t o t h e t ranspor t of gold in hydrothermal o re 2000 bars: Journal of Solution Chemistry, solutions: Geochirnica et Cosmochimica v. 4, p. 53-70. Ac ta , v. 37, p. 379-399.

(4) Ellis, A. J., 1963, Solubility of ca l c i t e in sodium (8) Ellis, A. J., and Giggenbach, W. F., 1971, chloride solutions a t high tempera tures : Hydrogen sulfide ionization and sulfur American Journal of Science, v. 261, hydrolysis in high t empera tu re solutions: p. 259-267. Ceochimica et Cosmochimica Acta , v. 35,

p. 247-260.

(9) Henley, R. W., Truesdell, A. K., and Barton, (10) Giggenbach, W. F., 1980, Geo the rma l gas P. B., 1984, Fluid minera l equilibria in equilibria: Geochimica e t Cosmochimica hydrothermal systems, Chapter 8: Ac ta , v. 44, p. 2021-2032. Reviews in Economic Geology, Society of Economic Geologists, v. 1.

Chapter 3 THE BEHAVIOR OF SILICA IN HYDROTHERMAL SOLUTIONS

Robert 0. Fournier

INTRODUCTION

Quar t z and chalcedony a r e t h e silica minerals commonly found in hydrothermal o re deposits. However, in many places the re is textura l evidence t h a t chalcedony formed a f t e r amorphous silica, probably with poorly crystall ine cristobalite or opal- C T a s an in termedia te phase. Fournier (1973) and apparently White (1965) used the t e rm B -cristobalite both for poorly crystall ine cristobalite and for opal- CT. Poorly crystall ine cristobalite shows broad X-r3y diffraction peaks centered a t about 4.1 and 2.5 A. Opal-CT also shows these same X-ray peaks plus an additional low-tridymite peak a t about 4.3 A (Jones and Segnit, 1971).

Quar t z is t he s table form of silica a t pressure- t empera tu re conditions found in convecting hydrothermal systems. Face ted qua r t z crys ta ls generally grow in solutions t h a t a r e not great ly supersa tura ted with silica, indicating relatively slowly changing conditions. In contras t , t h e deposition of atnorphous silica requires high degrees of si l ica supersaturation with respect t o quartz, and generally indicates large and rapid changes in the physical o r chemical na tu re of t he solution. These large and rapid changes may also a f f e c t t h e capacity of a solution t o t ranspor t and deposit ore. There a r e various ways t o bring about this supersaturation such a s rapid cooling (generally with decornpressional boiling), mixing of d i f ferent waters, pH changes, and react ion of t h e solution with volcanic glass. Each of these processes will be discussed in the subsequent sections. Much of what follows is taken from Fournier (1985).

SOLUBILITIES OF SILICA MINERALS

Experimentally determined solubilities of t h e common sil ica minerals in pure water a t t he vapor pressure of t h e solution up t o 300°C a r e shown in Figure 3.1. Equations fo r those portions of t hese curves in t h e t empera tu re range 0 - 2 5 0 ' ~ a r e given in the appendix. Equations expressing the molal solubility of quar tz and amorphous silica in water a t t he vapor pressure of t he solution over the t empera tu re ranges 20'-330°C and 90'-340'~ respectively a r e also given in the appendix, along with more general equations t h a t can be used t o ca lcula te qua r t z solubilities in water and saline solutions a t most t empera tu res and pressures of geologic in teres t . Increased pressure has l i t t l e e f f e c t a t 25OC, but causes significantly increased solubilities a s t h e t empera tu re is increased (Fournier and Po t t e r , 1982a). Going from

the vapor pressure of t h e solution t o 1000 bars increases t h e solubility of qua r t z about 19 percent a t 200°C, and about 36 percent a t 300 '~ .

Cr is tobal i te exhibits a wide range of crystall init ies (Murata and Nakata, 19741, and di f ferent samples may exhibit a complete spectrum of solubilities ranging from a-cr is tobal i te t o tha t of glass. Well-crystallized t r idymite should exhibit a solubility between those of a -cristobalite and chalcedony. When t w o o r more silica minerals a re in con tac t with a given solution, t h e most soluble silica phase will generally control aqueous sil ica until t h a t phase completely dissolves, is conver ted t o another more s table phase, or is taken ou t of con tac t with the circulating water by format ion - of intervening a l tera t ion products o r precipi ta tes (Fournier, 1973).

In t h e evaluation of hydrothermal processes, i t is of particular i n t e re s t t o determine conditions leading t o t h e precipitation of amorphous silica. Experimental determinations of t he solubility of amorphous silica in water a t high t empera tu res have been carr ied out by many investigators, including Fournier and Rowe (19771, and Marshall (1980). According t o Weres e t al.

Figure 3.1. Solubilities of various silica phases in water at the vapor pressure of the solution. A - amorphous silica, B = opal-CT (incorrectly identified as cl - cristobalite by Fournier, 1973), C = a - cristobalite, D = chalcedony,and E = quartz (from Fournier, 1973).

(1982), where degrees of si l ica supersaturation in respect t o amorphous sil ica reach a f ac to r of about 2.5, homogeneous nucleation is likely t o occur throughout a solution, si l ica polymers grow past cr i t ica l nucleus size, and colloidal par t ic les grow by fur ther deposition of si l ica on thei r surfaces. The colloidal par t ic les finally coagula te or flocculate, producing gelatinous ma te r i a l or f r iable aggregates of weakly cemen ted particles. Eventually, additional deposition of si l ica in t h e in ters t ices between these par t ic les may t ransform a friable aggregate t o relatively hard, dense material . If degrees of supersaturation a r e not g r e a t enough for homogeneous nucleation t o occur, deposit ion of dissolved sil ica takes place through heterogeneous nucleation, and growth of amorphous silica t akes p lace directly on solid surfaces. The resulting mater ia l is dense, vitreous sil ica t h a t initially should contain much less water than does gelatinous material . High salinity and near- neutra l pH favor more rapid r a t e s of amorphous silica deposition (Morey e t al., 1961; Makrides e t al., 1980; Cre ra r e t al., 1981; Weres e t al., 1982).

Natural hydrothermal waters range from very dilute t o concentra ted sa l t solutions. The e f f e c t s of dissolved sa l ts on t h e solubility of amorphous sil ica have been investigated by Marshall (19801, Marshall and Warakomski (19801, Marshall and Chen (1982a, 1982b), Chen and Marshall (19821, and Fournier and Marshall (1983). Below about 3 0 0 ' ~ most dissolved salts, with t h e exception of Na SO4, cause a slight decrease in the solubility o f amorphous silica. Solubilities of qua r t z and o the r si l ica minerals should be similarly a f f e c t e d by dissolved salts.

Addition of Na2S04 . great ly increases t h e solubility of amorphous sllica a t a l l measured temperatures (Chen and Marshall, 1982; Marshall and Chen, 1982a; Fournier and Marshall, 19831, apparently by the format ion of a silica-sulfate complex (Marshall and Chen, 1982b; Fournier and Marshall, 1983). It is not known if sodium is involved in t h a t complex, or if other dissolved sul fa tes also form complexes with silica. The format ion of a silica-sulfate complex also should increase t h e solubility of qua r t z in Na2S04-rich solutions. These silica-sulfate complexes may b e very important in some hydrothermal systems; a t Cesano, Italy, a hot, Na-K-SO4-rich brine with over 356,000 pprn to t a l dissolved s o l ~ d s a t 200' t o 300°C was encountered in a 1435-m deep geothermal well (Calamai e t al., 1976). Unfortunately, t h e concentration of dissolved silica in t h e brine a t Cesano is not well known.

THE BEHAVIOR OF DISSOLVED SILICA IN HOT-SPRING SYSTEMS

Much is known about t h e behavior of dissolved silica in ac t ive hydrothermal sys tems a s a result of studies of hot-spring waters (White e t al., 1956; Morey e t al., 1961; Fournier, 1973) and extensive drilling for geothermal resources (Mahon, 1966; Arnorsson and Sigurdsson, 1974; Arnorsson, 1975; Truesdell, 1976; Ellis and Mahon, 1977; Ellis, 1979; Henley and Ellis, 1983). In long-lived, presently ac t ive sys tems t h e

solubility of qua r t z has been found t o control dissolved silica in a l l geothermal reservoir waters a t t empera tu res greater t han about 180°c, most reservoir wa te r s above about 140°c, and many above 90°c. Chalcedony has a slightly higher solubility than quar tz (Fig. 3.1) and commonly controls silica a t t empera tu res below 90' t o 140°c, and somet imes a s high a s 1 8 0 ~ ~ . However, under special conditions for short periods of t ime, cr is tobal i te or volcanic glass may control dissolved silica, even a t very high temperatures . Special conditions include act ive faulting, hydrothermal explosions, and "short- circuiting" by drilling t h a t allows hot water t o come in to con tac t with glass- or cristobalite-rich rock t h a t has not previously been in con tac t with circulating hot water.

In ac t ive hydrothermal sys tems very l i t t l e silica appears t o precipi ta te underground when water rises relatively quickly from reservoirs where temperatures a r e less than 230' t o 250 '~ . This behavior allows t h e silica concentra t ion in a hot-spring or well water t o be used a s a chemical geothermometer t o e s t ima te underground reservoir t empera tu re (Fournier and Rowe, 1966; Mahon, 1966). Where reservoir t empera tu res a r e in excess of 230' t o 250°C, enough quar tz , chalcedony, or amorphous silica generally precipi ta tes f rom a n ascending solution t o cause ca lcula ted silica geothermometer temperatures applied t o hot-spring waters t o be significantly low. These empirically derived generalizations, based on field observations, a r e corroborated by experimental studies (Rimstidt and Barnes, 1980).

Self-sealing by deposition of amorphous silica a t t he tops of ac t ive geothermal sys tems has been discussed by Bodvarsson (19641, Facca and Tonani (19671, White et al. (1971), Coplen e t al. (19731, and Grindley and Browne (1976). Where quar tz controls dissolved sil ica a t depth, considerable cooling is generally required before amorphous silica can precipitate. Amorphous sil ica a t 1 0 0 ~ ~ has t h e s a m e solubility in pure water a s quar tz a t about 225OC. However, if an ascending solution rises f a s t enough t o cool adiabatically (through boiling in response t o lower pressure and without t ransfer of hea t t o t h e wallrock) and silica does not precipi ta te f rom solution during t h a t ascent , t h e sil ica concentration in the residual liquid will increase a s s t eam separa tes and a water in equilibrium with qua r t z a t a reservoir temperature of 210°C will yield a solution just s a tu ra t ed with respect t o amorphous sil ica a t IOOOC (Fig. 3.2).

In hot-spring sys tems a t Yellowstone National Park, Fournier and Truesdell (1970) found t h a t when silica does not precipi ta te from a thermal water during ascent f rom a n underground reservoir, t he re generally is an exchange of cations between t h e solution and t h e wallrock during upward flow. The reverse is t rue where significant amounts of silica precipi ta te from an ascending thermal water , l i t t l e or no cation exchange occurs. I t appears t h a t where a f r ac tu re is completely coated by precipi ta ted silica minerals, t h e wallrock may be effect ively removed from con tac t with the hydrothermal solution flowing through t h a t f rac ture . In t h a t event , ions in solution will not be able t o r e a c t with minerals in t h e wallrock in response t o changing temperature . Thus, t he precipitation of silica may

Temperature C0

Figure 3.2. Solubilities of amorphous silica and quartz (curve A) at the vapor pressures of the solutions. Curve B shows the amount of dissolved silica that would be present after adiabatic cooling (biling) to 10oOc, plotted as a function of the initial temperature of the solution before boiling. Point d shows the initial concentration of dissolved silica in equilibrium with quartz at 210°c, and pint e shows the concentration of dissolved silica after adiabatic cooling of that water from 210° to 100%.

adversely a f f e c t geochemical models applied to fluid flow t h a t assume continuous water-rock chemical equilibrium, or buffering of the pii of t he solution by reactions among s i l ica te minerals, unless those minerals precipi ta te along with quar tz , chalcedony, or amorphous silica. Where gelatinous silica coats t he walls of a vein, some react ion of the solution with t h e wallrock might continue a t a very slow r a t e by diffusion of ions through what amounts t o a chromatographic column.

In some presently ac t ive hydrothermal systems, such a s Broadlands, New Zealand, quar tz appears t o precipi ta te a t t he in t e r f ace where relatively saline, hot water from deep in the system mixes with overlying, less saline, colder water (W. A. J. Mahon, oral communication, 1982). Quar t z precipitation also may occur a t t h e cool margins of hydrothermal systems where hot and cold waters mix (Mahon e t al., 1980). The amount of si l ica supersaturation t h a t will occur a s a result of mixing waters having di f ferent initial temperatures can be determined graphically using an enthalpy-silica diagram (Fig. 3.3). In enthalpy-concentration diagrams, if hea t of mixing e f f ec t s a r e negligibly small, solutions t h a t result from

ENTHALPY ( Cal per g )

Figure 3.3. Solubility of quartz as a function of enthalpy at the vapor pressure of the solution. The straight lines A-C, E-C, and E-G show enthalpy-silica relations in solutions that result from mixing various proportions of hot and cold waters. See text for additional discussion.

mixing cold and hot components l ie along approximately s t ra ight lines joining t h e end members (Fournier, 19791, such a s line A-C in Figure 3.3. Cold groundwaters commonly have dissolved silica concentra t ions ranging from 10 t o 50 mg/kg. If t he solution a f t e r mixing has a relatively high enthalpy (high temperature) , such a s point D in Figure 3.3, qua r t z is likely t o precipitate. Note tha t even when two hot waters rnix, each in equilibrium with qua r t z a t d i f ferent temperatures , t h e resulting solutions may be supersa tura ted with respect t o quar tz (line: E-C in Fig. 3.3). However, mixing of some thermal waters, such a s E and G may give e i ther supersa tura ted (segment E- F) or undersa tura ted (segment F-G) silica solutions, depending on t h e proportion of t h e two waters in t h e mixture.

The cold-water component before mixing may be of relatively r ecen t me teo r i c origin and sti l l rich in oxygen f rom prior contact with the atmosphere. Mixing of th is oxygen-rich, meteor ic water with hot water could resul t in the formation of amethyst ine qua r t z because oxidizing conditions a r e required t o produce f e r r i c iron t h a t gives amethyst i t s purple color (Frondel, 1962).

Chalcedony might precipi ta te directly from some mixed waters a t appropr ia te low temperatures , but if t he enthalpy of t h e solution a f t e r mixing is 100 Cal/g, such a s point B in Figure 3.3, i t is likely t h a t l i t t le , if any, si l ica will deposit a s a resul t of t he mixing. This conclusion is based on t h e observation t h a t t h e sil ica mixing model of Truesdell and Fournier (1977) works very well when applied t o many hot-spring waters.

The thick, siliceous sinters t h a t a r e being deposited from presently ac t ive hot-spring sys tems a r e composed predominantly of amorphous mater ia l t h a t precipi ta tes f rom neu t ra l t o slightly alkaline waters. These waters a r e neutra l t o slightly acid in t h e reservoirs a t depth and commonly become alkaline during and a f t e r upward movement, owing t o loss of C 0 2 during boiling and evaporation. In contras t , highly ac idic waters (pH below about 3) tend not t o form thick siliceous deposits because hydrogen ions appear t o inhibit t h e polymerization of dissolved sil ica (Rothbaum e t al., 1979; Weres e t al., 1982). Hydrogen ions have been added t o some geothermal wa te r s t o prevent t he format ion of silica scale (Rothbaum e t al., 1979). From t h e above information, i t can be concluded t h a t th ick siliceous sinter deposits generally imply neutra l t o slightly alkaline wa te r s (alkaline by loss of C 0 2 ) t h a t have risen relatively quickly f rom underground reservoirs where the waters were in equilibrium with quar tz a t t empera tu res g rea te r t han about 210°C. Waters coming f rom reservoirs with t empera tu res a s low a s 160°C may become supersa tura ted in respect t o amorphous silica a s a result of cooling below 1 0 0 ~ ~ and a tmospher ic eva oration. However, silica t h a t precipi ta tes below 8 . 100 C IS likely t o be sof t , easily eroded dia tomaceous mud ra the r than hard sinter deposits t h a t a r e relatively resistant t o erosion.

I t i s not known what e f f e c t hydrogen ions might have upon precipitation of silica f rom Na-K-SO -rich brines. Because silica forms a complex with sut fa te , where sul fa te is removed from solution a s by format ion of alunite, exceedingly high-silica supersaturations may result , and silica may precipi ta te in an ac id environment.

Although boiling and cooling of slightly ac id t o slightly alkaline waters appears t o be t h e usual method of genera t ing silica solutions supersa tura ted with respect t o amorphous silica, o ther mechanisms a r e possible. These mechanisms involve very alkaline o r very acid waters, reactions with volcanic glass, or sudden decreases in fluid pressure.

ALKALINE WATERS

High pH is a possible, though unlikely, cause of high concentrations of dissolved silica in most natura l hydrothermal waters. Dissolved silica hydrolyzes t o form sil icic acid, H4Si04, and some of this si l icic ac id dissociates t o form H3Si04 (equation (0) in t h e Appendix). The amount of dissociation of si l icic ac id can b e calcula ted using equations (q), (r), and (s) t h a t a r e given in the Appendix. High pH favors increased dissociation and, therefore , increased amounts of dissolved silica in solution. Figures 3.1 and 3.2 show t h e solubilities of silica minerals in wa te r a t near-

neutra l t o acidic conditions. The e f f e c t of increased pH on t h e solubility of qua r t z a t t h e vapor pressure of t h e solution is shown in Figure 3.4. For comparison, t h e solubility of amorphous sil ica a t neutra l pH also is shown in Figure 3.4. A t t empera tu res above 130°C, qua r t z in alkaline solutions (pH 9.2 t o 9.5) is more soluble than is amorphous sil ica in neutra l solutions a t t h e s a m e temperature . Therefore , neutralization a t constant t empera tu re of an alkaline solution initially in equilibrium with qua r t z could lead t o precipitation of amorphous silica. This mechanism requires ( I ) t h a t a highly alkaline solution exis t a t high temperature , and (2) t h a t l i t t le o r no sil ica be consumed in the neutralization process.

Where water-rock chemical equilibrium is reached and quar tz plus mica or clay a r e present in the rock, hydrogen ions a r e buffered by react ions such a s

= K A ~ ~ ( A I S ~ ~ O ~ ~ ) ( O H ) ~ + 6Si02 + 2 ~ ' ( I ) Kmlca Quar t z

Quartz

0 0 100 2 00 300

Temperature, OC

F i g u r e 3.4. S o l u b i l i t y o f q u a r t z f r o m 25O t o 3 0 0 ~ ~ a t pH v a l u e s r a n g i n g f r o m 7 t o 1 0 ( s o l i d l i n e s ) . T h e d a s h e d l i n e s h o w s t h e s o l u b i l i t y o f a m o r p h o u s silica at n e u t r a l pH. Quartz s o l u b i l i t i e s a t n e u t r a l pH are from F o u r n i e r and P o t t e r (1982a), and i o n i - z a t i o n c o n s t a n t s f r o m F l e m i n g a n d C r e r a r ( 1 9 8 2 ) a n d S w e e t o n e t al. (1974) . Amor- phous silica s o l u b i l i t i e s a t n e u t r a l pH are from F o u r n i e r and Rowe (1977).

R. 0. FOURNIER 49

2.33NaA1Si206'H20 + 2H' Analc ime

+ S i 0 2 + 2.33H20 + 2 ~ a + (2) Quar tz

+ S i 0 2 + 2.33H20 + c a f 2 (3) Quar t z

and

The equilibrium constants fo r t h e s i l ica te pH-buffering react ions a r e such t h a t pH values a r e generally less t han 7, with higher salinity waters having lower pH values a t any given tempera ture , and wa te r s with given ca t ion ac t iv i t ies having lower pH values a t higher tempera tures . A t water-rock equilibrium conditions, pH values grea ter t han 7 a r e likely t o be found in hydrothermal solutions only where no phyllosilicates a r e present t o buffer t h e pH, o r where t h e r e is no normat ive or modal qua r t z in t h e rock. When t h e precipitation of quar tz cont ro ls aqueous si l ica concentrations, t h e ac t iv i ty of si l ica can be def ined equal t o unity so t h a t t h e equilibrium constant , Keq, fo r t h e react ion given by equation (1) is

with square brackets indicating activit ies. When dissolved silica is not controlled by qua r t z o r a more soluble silica mineral , such a s in u l t r amaf i c and nepheline-bearing rocks t h a t a r e devoid of si l ica minerals, t h e equilibrium constant is

and hydrogen ions decrease (pH increases) drastically in response t o t h e decreased silica activity. In near neut ra l and acid waters, dissolved si l ica appears t o be predominately H4Si04., With increasing pH, t h e solution is capable of d~ssolving increasing amounts of silica, but t h e ac t iv i ty of H4Si04 remains re la t ive ly low because much of t h a t dissolved s ~ l l c a is present a s H3SiO-.

I! t h e recharge water for a hydrothermal sys tem comes f rom a n alkaline, playa lake, continued water - rock react ion a t high t empera tu re (as recharge wa te r f lows through t h e sys tem) could eventually conver t a l l t h e pH-buffering clay and mica in t h e rock t o o the r minerals t h a t would allow sustained high pH values t o be r eached deep in t h e system. In t h a t even t , a n a lkal i

feldspar-carbonate assemblage would l ikely evolve and t h e isotopic composit ion of t h e o gen in t h e minerals would be relatively rich in 0 , ref lec t ing t h e evapora ted na tu re of t h e r echa rge water .

Subsurface boiling with removal of t h e evolved gas phase is another mechanism t h a t c a n gene ra t e high pH in a hydrothermal system. Chemical ly , th is is a non-equilibrium s i tua t ion in which continuing water - minera l reac t ions (such a s equat ions 1-4, proceeding f rom right t o le f t ) consume a lkal ies and l ibera te hydrogen ions a t a slower r a t e t han t h a t a t which hydroxyl i s genera ted by react ions such a s

mj = q + CH- ( 7 )

Down-hole wa te r samples col lec ted f rom research wells drilled in to ac t ive hot-spring sys t ems in Yellowstone National Pa rk (White e t al., 1975) generally have pH values of 6 t o 7, measured a t a i r t empera tu re immedia te ly a f t e r collection. However, one sample f rom t h e Y13 drill hole was found t o have a pH of 9.2 (R. Fournier, unpublished data). Tha t dri l l hole "short-circuited" a shallow permeable zone beneath t h e bot tom of t h e casing and a deeper permeable zone, separa ted f rom t h e shallow zone by impermeable rock. The d e e e r zone had a n init ial g t empera tu re of about 205 C , and was slightly overpressured re la t ive t o t h e weight of a n overlying column of wa te r everywhere a t boiling tempera ture . Because t h e shallow zone could a c c e p t wa te r f a s t e r t han t h e deep zone could supply i t , a s t eady s t a t e developed in which boiling occurred in t h e deep aquifer and s t eam flowed up t h e drill hole and o u t i n to t h e shallow aquifer, even when t h e wellhead valve was closed. As s t eam flowed f rom t h e sys tem, a n alkaline residual solution accumula t ed in t h e bot tom of t h e well, just below t h e deep permeable zone. A similar s i tua t ion might occur a s a resul t of ear thquake ac t iv i ty f rac tur ing t h e impermeable barrier. However, high gene ra t ez by evolution of acid-forming gases f rom a boiling wa te r is likely t o be short-lived and l imi ted in extent .

Where a high pH and a high dissolved si l ica concentra t ion appropr ia te fo r t h a t pH occur in a hydrothermal fluid, t h a t fluid must be neut ra l ized t o cause t h e precipitation of amorphous silica. Fur thermore , t h a t neut ra l iza t ion process must be relatively f a s t t o induce supersa tura t ion with respect t o quar tz , and i t mus t not consume a l a rge amount of si l ica re la t ive t o t h e amount of hydrogen ion produced. Most water-mineral reac t ions a r e likely t o proceed too slowly t o produce high deg rees of si l ica supersaturation. Also, reac t ions forming alkali feldspars, such a s equation (I) , consume t h r e e t imes a s much sil ica a s t h e amount of hydrogen ion t h a t is l iberated, and react ions involving montmoril lonite and analc ime o r wairaki te (equations 2 and 3) l ibera te t w o hydrogen ions for e a c h si l ica consumed. Anor th i te a l te r ing t o epidote l ibera tes hydrogen ion without consuming silica.

XaAI S i 0 + caf2 + 2H20 i n o ? t i i te

= Xa?p13SiP12(CH) + 2H+ (8) E p i d o t e

But , e x p e r i m e n t a l l y , e p i d o t e r e a c t s v e r y slowly, so q u a r t z would b e l ikely t o p r e c i p i t a t e if a so lu t ion w e r e n e u t r a l i z e d by t h i s reac t ion .

O t h e r m e c h a n i s m s f o r d e c r e a s i n g t h e a lka l in i ty of a h y d r o t h e r m a l so lu t ion i n c l u d e mix ing w i t h a n o t h e r n e u t r a l o r a c i d i c w a t e r , in f lux of a c i d g a s e s evolved f r o m a n o t h e r p a r t of t h e s y s t e m , ox ida t ion of sulf ides, a n d r e a c t i o n s involving n a t i v e sulfur .

ACID WATERS

Acid a l t e r a t i o n i s found in a n d a round m a n y o r e deposits . Q u a r t z d i sso lves a n d p r e c i p i t a t e s v e r y slowly in a c i d (pH <3) h y d r o t h e r m a l so lu t ions a t 200' t o 3 5 0 ' ~ using aqueous HCI a s t h e so lvent (R. Fourn ie r , unpublished da ta ) . Amorphous s i l i ca b e h a v e s sirni lar ly in HCI so lu t ions (I<amiya e t al., 1974; R o t h b a u m et al., 1979; Weres et al., 1982). T h e r e f o r e , w h e r e l a r g e a m o u n t s of a c i d a r e a d d e d t o a rock f r o m a n o u t s i d e source , o r quickly g e n e r a t e d wi th in a rock by oxida t ion reac t ions , a c i d a t t a c k upon fe ldspars and o t h e r s i l i ca tes m a y r e l e a s e s i l i ca t o solut ion, caus ing e x t r e m e s u p e r s a t u r a t i o n w i t h r e s p e c t t o q u a r t z and e v e n wi th r e s p e c t t o a m o r p h o u s si l ica. Neut ra l iz ing of s u c h a so lu t ion could c a u s e amorphous s i l i ca t o prec ip i ta te . T h e possibi l i ty t h a t s i l ica m i g h t p r e c i p i t a t e f r o m a n a c i d , su l fa te - r ich solut ion a s a consequence of f o r m a t i o n of a l u n i t e w a s d i scussed previously.

Acid i ty wi th in h y d r o t h e r m a l s y s t e m s c a n b e g e n e r a t e d in s e v e r a l ways. Acid-su l fa te a l t e r a t i o n is commonly o b s e r v e d a b o v e t h e w a t e r t a b l e in p r e s e n t l y a c t i v e hot-spring sys tems . T h e a c i d i t y resu l t s mainly f r o m oxida t ion of H2S t h a t m o v e s upward w i t h s t e a m t h a t h a s s e p a r a t e d f r o m underground boiling w a t e r , a n d p a r t l y f r o m oxida t ion of su l f ides in t h e rock. T h e s i l i c i f ica t ion t h a t a c c o m p a n i e s t h i s a c i d a t t a c k i s n o t d u e t o influx of s i l ica, b u t t o s t r o n g l e a c h i n g of a lka l ies t h a t a r e f lushed f r o m t h e r o c k by condensed s t e a m and m e t e o r i c w a t e r t h a t p e r c o l a t e down t o t h e w a t e r table.

Unt i l v e r y r e c e n t l y , dri l l ing f o r p roduct ion of g e o t h e r m a l e n e r g y i n a c t i v e h y d r o t h e r m a l s y s t e m s had e n c o u n t e r e d e x t r e m e l y a c i d condi t ions a t d e p t h a t only a f e w l o c a l i t i e s in z o n e s of a c t i v e vo lcan ism, s u c h a s Matsao in T a i w a n (Chen, 1970, 1975) a n d Onikobe (Yamada , 1976) a n d Matsukawa ( N a k a m u r a et al., 1970) in Japan . I t now a p p e a r s t h a t t h e r e i s deep , chloride-rich, a c i d t h e r m a l w a t e r a n d / o r d e e p a c i d a l t e r a t i o n (pyrophyl l i t e plus q u a r t z ) in m a n y a c t i v e h y d r o t h e r m a l s y s t e m s a s s o c i a t e d wi th a c t i v e or re la t ive ly young a n d e s i t i c volcanism. T h e s e inc lude Biliran (Lawless a n d Gonza les , 1982), Nasuji-Sogonon (Seas t res , 19821, Pa l impinon (Leach and Bogie, 1982), a n d Baslay-Dauin ( H a r p e r a n d Areva lo , 1982) in t h e Philippines, S u r e t i r n e a t (Heming e t al., 1982) in t h e New Hebrides, a n d Mirava l les in C o s t a R i c a (R. Fourn ie r , unpublished da ta ) . In s o m e of t h e s e s y s t e m s m u c h of t h e a c i d i t y m a y c o m e f r o m r e a c t i o n s involving m a g m a t i c gases , including S O H2S? a n d HCI.

A t M a t s a o a n d Onikobe acic?chlorlde w a t e r s wi th pH va lues l e s s t h a n 2 w e r e found a t d e p t h s g r e a t e r t h a n 1000 m e t e r s a t t e m p e r a t u r e s e x c e e d i n g 2 7 5 ' ~ . T h e reservoi r at M a t s a o i s in q u a r t z i t e , and a t Onikobe

i t i s in a n d e s i t e a l t e r e d t o pyrophyl l i t e a n d q u a r t z . W a t e r s c o l l e c t e d a t i n t e r m e d i a t e depths , and a l t e r a t i o n p r o d u c t s found in c u t t i n g s f r o m wells , show t h a t t h e s e deep , a c i d i c w a t e r s a r e n e u t r a l i z e d by mix ing wi th shallow ground w a t e r a n d by r e a c t i o n wi th over ly ing v o l c a n i c r o c k s a s t h e w a t e r s r i se t o w a r d t h e s u r f a c e . T h e r e f o r e , t h e a c i d i t y d o e s n o t a p p e a r t o b e t h e r e s u l t of d o w n w a r d m o v e m e n t of d i rec t ly over ly ing w a t e r s t h a t h a d b e c o m e a c i d i c by s u r f a c e oxidation. Ell is (1977) a t t r i b u t e d t h e d e e p a c i d i t y a t Matsao t o t h e r e a c t i o n of w a t e r w i t h d e e p l y buried n a t i v e su l fur depos i t s , p roducing s u l f u r i c a c i d a n d hydrogen sulf ide,

However , a s mentionGd above , s o m e o r a l l of t h e a c i d i t y could r e s u l t f r o m i n t e r a c t i o n of w a t e r wi th g a s e s evolved f r o m a c rys ta l l i z ing m a g m a a t dep th , o r f r o m hydrolysis r e a c t i o n s b e t w e e n s a l t a n d w a t e r t h a t o c c u r a t high t e m p e r a t u r e s a n d low pressures.

Iwasaki a n d O z a w a (1960) a n d Saki a n d M a t s u b a y a (1977) p r e s e n t e v i d e n c e f o r t h e g e n e r a t i o n of a c i d i t y by t h e r e a c t i o n

I t is l ikely t h a t s u l f u r i c a c i d a l s o c a n b e g e n e r a t e d by o t h e r mechanisms . T h e 1982 e r u p t i o n of El Chichon volcano in Mexico c o n t r i b u t e d f a r m o r e su l fur ic a c i d t o t h e a t m o s p h e r e a n d s t r a t o s p h e r e t h a n i s usual f o r cornparab ly s i z e d e r u p t i o n s of o t h e r volcanoes, s u c h a s Mount St . Helens (B. Toon, o r a l communica t ion , 1982). I t i s n o t e w o r t h y t h a t gypsum beds o c c u r in t h e s e d i m e n t a r y s e c t i o n b e n e a t h E l Chichon , b u t a r e n o t p r e s e n t b e n e a t h Mount St . Helens. Apparen t ly r e l a t i v e l y oxidized, sulfur-r ich g a s e s m a y b e evolved w h e r e gypsum o r a n h y d r i t e a r e involved in hydrolysis r e a c t i o n s a t high t e m p e r a t u r e s

T h e C a O t h a t f o r m s by r e a c t i o n (11) will r e a c t w i t h q u a r t z a n d / o r o t h e r s i l i c a t e s t o f o r m a v a r i e t y of c a l c i u m s i l i ca tes . T h e i s o t o p i c compos i t ion of su l fur in t h e s e v o l c a n i c s y s t e m s should b e usefu l f o r dist inguishing b e t w e e n s u l f a t e der ived f r o m gypsum a n d s u l f a t e d e r i v e d f r o m volcanic SO2.

T h e i m p o r t a n c e of HCI a s a c a u s e of a c i d i t y in h y d r o t h e r m a l s y s t e m s should n o t b e overlooked. O v e r 7000 m g / k g C1 a s HCI w e r e found in d r y s t e a m c o m i n g f r o m a shallow wel l d r i l l ed a t H a k o n e volcano in J a p a n (Kimio Noguchi , o r a l c o m m u n i c a t i o n , 1970). S o m e o r a l l of t h a t HCI m a y h a v e been g e n e r a t e d b y hydrolysis of N a C l at m o d e r a t e t o high t e m p e r a t u r e s a n d low pressures

NaCl + H20 = NaCH + E l ( 1 2 )

Many i n v e s t i g a t o r s h a v e found HCI in c o n d e n s a t e a f t e r c i r c u l a t i n g d r y s t e a m o v e r sol id NaCl (Briner a n d R o t h , 1948; M a r t y n o v a a n d Samoilov, 1959; Galobardes e t al., 1981). In e x p e r i m e n t s at 6 0 0 ' ~ I found t h a t s ign i f ican t a m o u n t s of HCI a r e g e n e r a t e d by r e a c t i o n (12) a t p ressures be low a b o u t 350 bars, wi th m o r e HCI produced at l o w e r pressures. Addit ion of q u a r t z t o t h e

sys tem great ly increased the yield of HCI. This occurs because NaOH i s removed from the solution by react ion with quartz, with precipitated sodium s i l ica tes a s products. Solubilities of sodium sil icates decrease with increasing temperature (Rowe e t al., 1967). In na tu ra l systems, where aluminum is available in clays and o the r minerals, albit ization is likely t o resul t f rom t h e hydrolysis of NaCI.

In some places, deep acidity may result from downward movement of wa te r t h a t has become acid a t and near t h e wa te r table (Oki and Hirano, 1970; Henley and Ellis, 1983). In the Norris Geyser Basin region of Yellowstone National Park, acid waters a r e generated high on a hillside where H2S is oxidized t o sulfate. Some of this wa te r appears t o percolate hundreds of me te r s underground where i t mixes with high- t empera tu re (%27o0c), neutra l water rich in chloride. The resulting "acid-chloride-sulfate" waters, t h a t issue a s hot springs and geysers, have been extensively analyzed (Gooch and Whitfield, 1888; Allen and Day, 1935; Rowe e t al., 1973).

REACTION WITH GLASS

I t was noted above t h a t t h e solubility of glass can control dissolved silica concentrations where new f rac tu res bring hot water in to con tac t with previously unaltered rock. Laboratory experiments a t hydrothermal conditions show t h a t volcanic glass contributes si l ica t o solution t o about t h e same ex ten t a s pure amorphous sili,ca (Dickson and Po t t e r , 1982). In the absence of rapid precipitation of quartz, chalcedony, o r cristobalite, any cooling of a solution t h a t has r eac ted extensively with volcanic glass could result in t h e precipitation of amorphous silica. The f i rs t s t age in t h e precipitation process is likely t o be polymerization of t h e dissolved silica, with formation of colloidal and gelatinous par t ic les t h a t can be swept along in a moving fluid.

React ion of hot water with previously unaltered glass-rich volcanic rock may explain exceedingly high silica concentrations found in one vigorously discharging, boiling hot spring in Yellowstone National Park t h a t is presently depositing gold-bearing sinter. Tha t spring f i rs t appeared shortly a f t e r t he magnitude 7.1 Hebgen Lake ear thquake in 1959. When I f i rs t observed and sampled this hot spring, i t was discharging more than 300 Ilmin., t he water was fountaining a s much a s a me te r above t h e vent, and i t had an opalescent appearance owing t o dispersed colloidal silica. The deposition of siliceous sinter around t h e edge of the pool and along the channels where water overflowed was especially rapid where water c a m e in con tac t with organic material . One sample of this s in ter was analyzed for gold by neutron activation; i t contained 0.8 mglkg. In this instance the gold could have been co-precipitated with t h e colloidal silica a t depth and then transported t o the surface along with t h a t silica.

AMORPHOUS SILICA-CHALCEDONY RELATIONS

Recent ly deposited hot-spring siliceous sinters have been found t o be composed predominantly of

amorphous material . Sinters a f ew thousand years old tend t o have an opaline appearance, especially if buried by younger deposits, and e i ther a poorly crystall ine cr is tobal i te or opal-CT X-ray pattern. According t o T. E. C. Kei th (wri t ten communication, 19821, many X-ray di f f rac t ion pat terns of older s in ters f rom Yellowstone National Park indicate poorly crystall ine cr is tobal i te without the small t r idymite peak t h a t is cha rac te r i s t i c of opal-CT. In contras t , t h e sil ica phase l e f t a s a residue in acid-altered and leached environments in Yellowstone National Park generally does exhibit a broad 4.3 A tr idylnite X-ray peak in addition t o t h e broad cristobalite peaks.

Many s in ters t ens of thousands of years old a r e chalcedonic, especially if buried and exposed t o higher t empera tu res than those generally a t ta ined a t t h e ear th ' s surface. The transformation of amorphous sil ica t o poorly crystall ine cristobalite, opal CT, chalcedony, and qua r t z has been studied in t h e field and experimentally by many investigators (White e t al., 1956; C a r r and Fyfe, 1958; Fyfe and McKay, 1962; Heydemann, 1964; Ernst and Calver t , 1969; Mizutani, 1970; Murata and Nakata, 1974; Murata and Larson, 1975; Murata and Randall, 1975; Kastner e t al., 1977; Hein and Scholl, 1978; Keith and Muffler, 1978; Kano and Taguchi, 1982). Time, high t empera tu re , high pH, high salinity, and the presence of dissolved Mg al l have been found t o favor the t ransformat ion of amorphous silica to quar tz (or chalcedony), with e i the r poorly crystall ine cristobalite or opal CT as a n in termedia te phase. I t is possible t h a t o ther dissolved consti tuents also may cata lyze t h e transformation.

Apparently, chalcedony may form ei ther by d i r ec t precipitation from solution without going through an amorphous silica s tage , or by transformation of amorphous silica to crystall ine material . I t is important t o know if chalcedony formed a s a primary precipi ta te or by transformation of amorphous silica because of the implications about conditions required t o precipi ta te one or t h e other. Morphologic fea tures , such a s dehydration cracks, slump s t ructures , and thicker silica deposits on t h e bot toms of cavi t ies than on the sides and tops (gravitational se t t l ing of amorphous silica particles) indicate t h a t amorphous sil ica was present initially. In t h e absence of such features , t he mechanism of format ion of a given chalcedony is in doubt. The t race-e lement contents of primary chalcedonies and those formed by transformation of amorphous silica should be compared to determine if t he re a r e diagnostic differences. Elements t h a t form small, highly charged, cations, such a s Ge, Ga, B, and Al, might subst i tu te for Si more readily in amorphous mater ia l than in crystall ine and, therefore, might be useful in distinguishing among types of chalcedony.

SPECULATIONS REGARDING SOME TEXTURES OF QUARTZ

Where qua r t z precipitates in open spaces directly frorn hydrothermal solutions i t exhibits crys ta l faces, and all t h e crystallographic directions generally a r e roughly perpendicular t o a surface upon which growth

initially occurred. However, many hydrothermal veins exhibit bands of randomly or iented equant grains of, or uniform filling by, anhedral quar tz t h a t show no indications of ear l ier euhedral s tages of growth. In deep deposits t h e r e t end t o b e no cavi t ies in these veins. In shallower deposits some cavi t ies may be present, and portions of qua r t z grains t h a t extend in to cavi t ies a r e faceted. These textures suggest t h a t t h e qua r t z may have formed by growth within a band t h a t was initially amorphous silica, or chalcedony (less likely). If so, fluid inclusions within tha t qua r t z e i the r were introduced a f t e r crystall ization, or represent fluid t h a t was t rapped a t t h e t i m e of format ion of t h e amorphous or chalcedonic silica and la ter remobilized during t h e t ransformat ion t o quartz. Therefore, fluid- inclusion filling t empera tu res and salinit ies may give l i t t le or no information about t h e condition of t h e hydrothermal fluid t h a t initially deposited t h e siliceous vein material .

The s ize and general appearance of qua r t z grains a f t e r amorphous silica probably depends upon many factors , including the nucleation mechanism and init ial water con ten t of t he amorphous silica, t he pH and chemical composition of t h e surrounding fluid, t h e t empera tu re a t which the amorphous silica deposited, and t h e r a t e a t which t h e system cooled. In general, higher t empera tu res of transit ion of amorphous sil ica t o qua r t z will result in coarser grained quartz. Equant, anhedral grains of milky-looking quar tz with many tiny fluid inclusions might indicate crystall ization from gelatinous amorphous sil ica t h a t formed by homogeneous nucleation f rom a solution very supersa tura ted with silica. A likely mechanism for genera t ing highly supersaturated silica a t high t empera tu res is by a sudden decrease in fluid pressure (Fig. 3.5); possibly from nearly l i thos ta t ic t o hydrosta t ic or lower. Grindley and Browne (1976) a t t r ibu ted the format ion of strongly silicified breccias adjacent t o open fissures a t Wairakei and Broadlands, New Zealand t o sudden decreases in fluid pressure resulting from hydraulic f rac tures propagating through self-sealed rock. The silicifying minerals a r e qua r t z plus adularia, usually accompanied by pyrite. Kei th and Muffler (1978) suggested t h a t simultaneous brecciation and deposition of amorphous silica occurred a s a result of a sudden decrease in pressure, caused e i ther by f rac tur ing accompanying resurgent doming, or draining of a glacial lake t h a t decreased t h e local hydrostatic pressure. Self-sealing and e f f e c t s of sudden decreases in fluid pressure will be discussed fur ther in the section on quar tz solubility a t high temperatures .

Jasperoid and Massive Replacement of Limestone by - . Silica -

The t e rm jasperoid is of ten used t o ca tegor ize bodies of massive silica (commonly containing iron sulfides o r oxides), irrespective of t he size, shape, internal s t ructure , or geologic setting. Lovering (1972) has described their character is t ics and distribution in t h e United States. I t is likely tha t they form by many mechanisms. Some may be recrystall ized hot-spring sinter deposits. Others may have formed where rising, hot wa te r mixed with shallow, cold water. Sti l l o the r s

may have formed as a result of vigorous decompressional boiling, particularly where overpressured thermal fluid expanded in to open, hydrostatically pressured cavities. Such throt t l ing processes a r e discussed by Barton and Toulmin (1961) and Toulmin and Clark (1967). Most jasperoids, however, appear t o be massive sil ica replacements of l imestone or, much less commonly, dolomite (Lovering, 1972). They a r e particularly prevalent in carbonate- hosted gold deposits, such a s Carlin (Radtke et al., 1980; Rye, 1985) and those in t h e J e r r i t t Canyon Dis t r ic t (Hawkins, 19821, where massive, fine-grained sil ica (now quar tz) replaces l imestone with l i t t l e or no associated calc-silicate. In those deposits the silicification of lirnestone is most pronounced near faul ts and where the l imestone was initially most permeable.

Silica replacement of l imestone requires the simultaneous dissolution of ca l c i t e and precipitation of silica. Below 300°C, a t a constant par t ia l pressure of COZ, ca lc i te becomes more soluble with decreasing t empera tu re (Ellis, 1959, 1963; Holland and Malinin, 19791, while t h e solubilities of quar tz , chalcedony, and amorphous silica decrease (Fig. 3.1). Also, a t constant temperature , ca l c i t e becomes less soluble a s the par t ia l pressure of C 0 2 decreases (Ellis, 1959). Therefore, slow cooling ( w ~ t h o u t boiling) of a solution with a near neutra l pH should promote replacement of l imestone by silica. The sil ica phase t h a t precipi ta tes i s likely t o be quar tz or chalcedony, because of t he slow cooling. I t is less likely t h a t l imestone will be replaced by sil ica where a near-neutral solution boils. Decompressional boiling results in loss of dissolved C O a s well a s a rapid dec rease in temperature , and 2? c a l c ~ t e may dissolve or precipitate, depending on t h e composition of t h e hydrothermal solution, t h e pressure a t which boiling is init iated, t he drop in pressure, and whether t h e system is open or closed t o loss of volati les during t h a t boiling. Escape of volati les from t h e system during boiling is particularly likely t o cause deposition of ca lc i te , which would not favor replacement of l imestone by silica.

I t rnay be possible t o accomplish massive replacement of l imestone by silica with l i t t l e or no cooling if t he hydrothermal solution is acidic, even when t h a t solution is just s a tu ra t ed with silica in respect t o quartz. React ion of a silica-saturated, ac idic solution with l imestone will genera te C 0 2 , and some or a l l of i t will dissolve in t h e solution, depending on severa l fac tors , including t h e temperature , partial pressure of C 0 2 , pH, and salinity. Dissolved C 0 2 lowers t h e ac t iv i ty of wa te r and decreases t h e solubility of quar tz (Shettel, 1974). Therefore, quar tz should precipi ta te a t t he in t e r f ace of solution and limestone where C 0 2 is generated. If t h e acidic solution were initially supersa tura ted with silica with respect t o quar tz , increasing t h e pH of t h e solution by reaction with ca lc i te would probably cause precipitation of silica, even without t h e C 0 2 e f fec t . The deposited silica might be quartz, chalcedony, or amorphous, depending on t h e init ial silica concentration in t h e acid solution and t empera tu re a t which the solution r eac t s with t h e limestone. By t h e above C 0 2 mechanism, sil ica replacement of l imestone a t constant t empera tu re would be more

likely deeper in t h e system where fluid pressures a r e 5000

high enough for significant amounts of C O t o dissolve in t h e solution a t t h e point where ?hat gas is generated. A t shallow levels, where t h e to t a l fluid pressure i s too low t o allow much C 0 2 t o dissolve in the solution, si l ica replacement of l imestone is more likely t o occur a s a result of reaction of t h a t l imestone with an acid solution t h a t has become supersa tura ted 2 with sil ica a s a result of cooling, e i ther by boiling or > 3000 conductive h e a t loss. E

Quar tz Solubility a t High Temperatures

I t was previously noted tha t below about 300°c, pressure has a moderate a f f e c t and added sa l t has l i t t le a f f e c t upon t h e solubility of quartz. Above 300°c, both pressure and added sa l t a r e very important. Calcula ted solubilities of quar tz in pure water over a wide range of t empera tu res and pressures, using t h e equation of Fournier and Po t t e r (1982a), a r e shown in Figure 3.5. In t h a t f igure the re is a solubility maximum (reported by Kennedy, 1950) t h a t extends from about 3 4 0 ' ~ a t t h e vapor pressure of solution t o 520°C close t o 900 bars. The shaded a r e a in Figure 3.5 shows a region of re t rograde solubility in pressure-temperature space. Where cold water is hea ted a t constant pressure less than about 900 bars, i t will dissolve more and more silica until e i ther t he solution s t a r t s t o boil (a t pressures below about 165 bars) or t he solubility maximum is reached. With fur ther heating tha t water will precipi ta te quartz. The precipitation of quar tz in deep par ts of a hydrothermal system may decrease the permeabili ty t o such a n e x t e n t t h a t l i t t le convecting me teo r i c wa te r can a t t a i n temperatures much greater than those shown by the quar tz solubility maximum in Figure 3.5 (Fournier, 1977, 1983a, 1983b). However, computer modeling by L. A. Keith and P. T. Delaney (wri t ten communication, 1984) shows tha t a completely impermeable seal is not likely t o result in rea l is t ic t imes solely by deposition of quar tz (or a mixture of qua r t z and other minerals) from a solution t h a t is hea ted a s i t flows toward a hea t source. This i s because the process is self-limiting: a s permeabili ty is decreased by quar tz deposition, the r a t e of flow decreases, which, in turn, decreases t h e r a t e of t ranspor t of silica to the place where deposition can occur. However, o ther fac tors also may contr ibute t o t h e a t t a inmen t of a completely impermaeble seal, such a s quasi-plastic flow of rock t h a t takes p lace a t increasingly rapid r a t e s a s temperature increases. Tempera tu re profiles calculated from heat-flow d a t a for severa l locali t ies in the western United Sta tes , and ear thquake focal depths a t those same locations, show t h a t t h e temperature a t which deformation changes f rom frictional (br i t t le f rac ture) t o quasi-plastic flow ranges f rom about 300' t o 4 5 0 ' ~ (Smith and Bruhn, 1984). This overlaps the 350' t o 5 0 0 ' ~ t empera tu re range in which self sealing by precipitation of qua r t z and o the r minerals is likely t o occur when solutions a r e hea ted a t constant pressure (Fournier, 1977, 1983a, 1983b; Sleep, 1983). Therefore, because of t h e above mentioned permeabili ty reduction processes, t he t i m e in terval over which meteor ic water a t hydrosta t ic pressure may in t e rac t directly with a shallow int ruded

Quartz solubility in w a t e r

T e m p e r a t u r e , C

F i g u r e 3.5. C a l c u l a t e d s o l u b i l i t i e s o f q u a r t z i n w a t e r up to 9 0 0 ~ ~ a t the i n d i c a t e d p res - s u r e s . The shaded area emphas izes a r e g i o n o f r e t r o g r a d e s o l u b i l i t y .

body of magma (or st i l l very hot rock) may be l imited t o t h e early s t age of development of t h e hydrothermal system, or episodically the rea f t e r with creat ion of new f rac tu res by t ec ton ic ac t iv i ty or thermal o r hydraulic cracking (Secor, 1965; Phillips, 1973; Henley and McNabb, 1978).

This model has some in teres t ing implications f rom t h e point of view of o re genesis. With the circulation of meteor ic water through shallow, intrusive rocks c u t off a t a n ear ly s t age in t h e cooling process, those intrusive rocks will cool a t a slower r a t e , and hydrothermal ac t iv i ty will continue for a longer t ime. This i s because hea t must be t ransferred by conduction from t h e remaining very hot rock t o t h e cooler convecting hot water. Conductive t ransfer of thermal energy is much less ef f ic ient than convective transfer.

Hydrothermal explosion act iv i ty is another possible consequence of t h e deposition of an impermeable qua r t z seal (Henley and McNabb, 1978). Large and s t eep t empera tu re and pore-pressure gradients a r e likely t o evolve where an impermeable zone becomes established about a hea t source. Even though convective flow of me teo r i c wa te r is cu t off from t h e outside, t h e pore spaces within the zone between t h e quartz-sealed barrier and t h e remaining very hot rock a r e likely t o contain wa te r or brine. This fluid may be a l l or par t me teo r i c o r connate water, l e f t over from before t h e silica sealing became complete. However, some or a l l of t h a t fluid could be volati les evolved from a crystall izing magma (Burnham, 1967, 1979). If volati les continue t o be evolved from a crystall izing magma, i t i s easy to envision a si tuation in which t h e fluid pressure on the high-temperature side of t h e qua r t z seal becomes very

la rge (Phillips, 1973); suff ic ient ly la rge t o cause format ion of a b recc i a pipe o r even a conduit fo r a volcanic erupt ion (Morey, 1922).

Hydraulic f r ac tu r ing will occur when t h e pore- fluid pressure exceeds t h e confining pressure ( the l ea s t principle s t ress) by a n a m o u n t equal t o t h e tensile s t r eng th of t h e rock. The confining pressure may range f rom less t han normal hydros ta t ic t o l i thos ta t ic , depending on whether open f i ssures a r e present, t h e na tu re of t h e fluid in those fissures, and permeabili ty relations. Propogation of e i t he r a hydraulic o r t ec ton ic f r a c t u r e through impermeab le rock f rom a region of high fluid pressure in to a region of lower fluid pressure may cause a significant decompression of t h e high-pressure fluid. If t h e t he rma l energy in t h e decompressing liquid and surrounding rock is large, massive flashing of wa te r t o s t e a m may result. The expanding s t eam may explosively propel rock f r agmen t s i n to t h e a l r where flashing occurs a t relatively shallow levels, and in to cavi t ies and open fissures a t deeper levels. Even without a magmat i c contribution t o t h e t rapped fluids, pore pressures of those fluids could increase sufficiently t o rupture t h e enclosing rock a s a resul t of conduct ive heating.

Hydrothermal explosive ac t iv i ty may be a n impor t an t cont r ibut ing f a c t o r t o o r e deposition fo r various reasons, both physical and chemical. Brecciation grea t ly increases t h e permeabili ty, providing easy a c c e s s f o r l a t e r hydrothermal f luids t h a t may deposit ores. When a hydrothermal explosion occurs, a lo t of wa te r is conver ted t o s team. Other volati le cons t i tuents , such a s H2S and C O , initially dissolved in t h e liquid phase, a r e parti t ioned in to t h a t s team. This parti t ioning of volatiles, in turn, m a y inc rease t h e pH of t h e residual liquid. A t t h e s a m e t i m e t h e concentra t ions of t h e non-volatile cons t i tuents remaining in t h e liquid phase increase a s a resul t of t h e separa t ion of s t eam, while t h e solubilities of minerals generally dec rease a s a resul t of t h e dec rease in pressure. In Figure 3.5, no t e t h e large decrease in qua r t z solubility with decreasing pressure a t t empera tu re s above about 340 '~ . In addit ion, where boiling occurs, t h e t empera tu re of t h e sys tem should dec rease because the rma l energy is required t o conver t liquid wa te r t o s team. The above f ac to r s generally favor deposit ion of silica, sulfides, and noble metals. Whether o r not o r e is deposited will depend in pa r t on t h e me ta l and sulfur content of t h e init ial fluid. However, when init ial t empera tu re s a r e above about 340°c, qua r t z and o the r minerals should prec ip i ta te when and where t h e r e is a sudden drop in pore pressure. Also, any K-feldspar t h a t prec ip i ta tes along with t h e qua r t z a s a resul t of a sudden drop in pressure is likely t o be more potassium-rich than t h a t which was in equilibrium with t h e fluid prior t o t h e drop in pressure (Fournier, 1976). A t Wairakei and Broadlands, quar tz , adularia, and generally pyr i te a r e t h e phases observed in t h e hydrothermal breccias t h a t formed a t about 2 0 0 ~ - 3 0 0 ~ ~ (Grindley and Browne, 1976). Where exceptionally high degrees of si l ica supersaturation occur , part icularly a t lower tempera tures , amorphous si l ica may prec ip i ta te and then a l t e r t o qua r t z o r chalcedonic silica.

Many of t h e conclusions in t h e above discussion were based on t h e solubility behavior of quar tz in pure

water. The e f f e c t s of added sa l t s can be modeled using NaCl solutions. Ca lcu la t ed solubilities of qua r t z in aqueous NaCl, using t h e method of Fournier (1983b), show t h a t adding dissolved sa l t s should change t h e positions of t h e qua r t z solubility maxima and t h e e x t e n t of t h e field of r e t rog rade qua r t z solubility t h a t a r e shown in Figure 3.5. However, t h e conclusion t h a t qua r t z deposit ion should contr ibute significantly t o t h e format ion of a n impermeable barr ier t h a t prevents fluids a t hydros ta t ic pressure f rom in terac t ing directly with very hot rock o r magma , is not changed by adding sa l t t o t h e system. Figure 3.6 shows t h e e f f e c t of 5 and 18 weight percent NaCl upon qua r t z solubility a t 500 bars pressure and high tempera tures . Added NaCl grea t ly increases t h e solubility of qua r t z a t t empera tu re s above about 300°c, and shi f t s t h e solubility maximum toward higher tempera tures . A t very high concentra t ions of sa l t , t h e vapor-pressure curve may be in t e r sec t ed before a solubility maximum is a t t a ined , such a s a t point A in Figure 3.6. In t h a t event , t h e solution will boil if t h e t empera tu re is increased fu r the r without increasing t h e pressure. Wherever a solution boils, t h e concentra t ion of dissolved si l ica in t h e residual solution should increase a t a more rapid r a t e t han c a n be accommodated by t h e increasing salinity. T h e amoun t of silica t h a t can dissolve in t h e newly fo rmed gas o r s t eam phase is generally insufficient t o o f f se t t h e supersa tura t ion genera ted in t h e residual brine, and qua r t z should prec ip i ta te because t h e t empera tu re is high. Thus, qua r t z veining should occur e i ther when a d i lu te solution is hea t ed t o a t empera tu re above t h e qua r t z

Quartz solubi l i ty i n NaCl solutions /-. 500 bars

/ /' A

/ /

/ /

8 . /

240 I I I

300 400 500 Temperature, OC

Figure 3.6. Comparison of calculated quartz, solubilities (Fournier, 198313) in water and 5 and 18 weight percent aqueous NaCl at 500 bars and the indicated temperatures.

solubility maximum o r when a saline solution exceeds t h e t e m p e r a t u r e of t h e vapor-pressure cu rve a t a given pressure.

Tempera tu re s and approximate depths a t which boiling will occur in wa te r and 5 and 18 weight pe rcen t NaCl solutions a r e shown in Figure 3.7. Hydros ta t ic conditions a r e assumed in Figure 3.7, with pressure fixed by a n overlying liquid with a n ave rage density of I throughout a ver t ica l column up t o t h e ground surface. Re la t ive t o t h e depth scale, lower assumed ave rage densit ies will move t h e boiling point curves down, and higher assumed densit ies will move t h e curves up. The approximate positions of t h e qua r t z solubility maxima f o r wa te r and 5 weight percent NaCl also a r e shown in Figure 3.7.

Because t h e init ial permeabili ty severa l k i lometers deep in a hydrothermal sys tem is likely t o be l imi ted t o a f ew widely spaced f r ac tu re s o r f r ac tu red zones of rock, a n irnpermeable zone result ing in g r e a t par t from qua r t z deposition in those f e w f r a c t u r e s may go unrecognized a s a significant fea ture . Also, in fossil hydrothermal sys tems where e s t ima ted t empera tu re s a t t h e t i m e of vein format ion a r e g rea t e r t han 340°c, i t may be difficult t o de t e rmine whether a given qua r t z vein deposited a s a resul t of increasing o r decreas ing tempera ture . If t h e r e is o ther hydrothermal a l te ra t ion associa ted with t h e qua r t z deposition, t h a t a l t e r a t ion may give a n indication of t h e t he rma l history: a lb i te is likely t o form in veins and a f t e r K-feldspar where a solution is heat ing and K-feldspar or ~ n u s c o v i t e would be

Figure 3.7. Temperature vs. depth (pressure) diagrams showing boiling point curves for pure water (curve A), 5 weight percent aqueous NaCl (curve B), and 18 weight percent aqueous NaCl (curve C). Also shown are the positions of quartz solubility maxima in water (curve D) and 5 weight percent aqueous NaCl (curve E).

deposited in veins and a f t e r plagioclase where a solution is cooling (Hemley e t al., 1971).

CONCLUSIONS

In well-established hydrothermal sys tems, where wa te r remains in con tac t with t h e surrounding rock a t a given high ternpera ture f o r more than a f e w days or weeks, qua r t z cont ro ls aqueous si l ica (Rimstidt and Barnes, 1980). Slow cooling of a hydrothermal solution generally will resul t in t h e deposition of qua r t z if init ial t empera tu re s a r e be tween about 200' and 340°C. Rapid cooling allows supersa tura ted si l ica solutions t o form, particularly when t h e cooling is predominantly t h e resul t of decompressional boiling. Supersa tura ted si l ica solutions a lso may evolve where hot wa te r dissolves glass-bearing rocks, and where rocks a r e a l t e r ed by very ac id solutions. High alkalinity (high pH) i s generally not impor tant in mos t na tura l hydrothermal sys tems, but might be a f ac to r in a few places fo r shor t periods of time.

Slight si l ica supersa tura t ion in respect t o qua r t z is required fo r chalcedony t o prec ip i ta te d i rec t ly f rom solution. Chalcedony appea r s t o form and persist only a t t empera tu re s below about 180°C. Large degrees of si l ica supersa tura t ion a r e required fo r amorphous si l ica t o prec ip i ta te . Voluminous deposits of siliceous s in ter generally indica te deposit ion f rom neutra l t o slightly alkaline (by loss of C02) , chloride-rich wa te r s t h a t flowed quickly t o t h e su r f ace f rom a reservoir with a t empera tu re in t h e range 200' t o 270°C. Waters flowing f rom reservoirs with lower t empera tu re s contain t o o l i t t l e si l ica t o form thick, hard, s in ter deposits. Waters flowing f rom reservoirs with t empera tu re s above 2 7 0 ' ~ conta in so much dissolved si l ica t h a t significant amounts prec ip i ta te in t h e channelways leading t o t h e surface , stopping hot- spring ac t iv i ty be fo re large s in ter deposits can form. Very l i t t l e si l ica prec ip i ta tes f rom waters with pH values below about 3 t o 4.

Amorphous si l ica i s relatively unstable and t ransforms t o poorly crys ta l l ine cristobali te, opal C T and chalcedony o r quartz. The t i m e required fo r t hese t ransformat ions depends upon t empera tu re and t h e composit ion of f luid in con tac t with t h e amorphous material . Morphologic f e a t u r e s show t h a t some chalcedonies have formed a f t e r amorphous silica. Where such morphologic f ea tu re s a r e absent , t h e origin of a given chalcedony is in doubt. I t is impor tant t o know if chalcedony fo rmed a s a primary prec ip i ta te o r by t ransformat ion of amorphous si l ica because of t h e implications about conditions required t o prec ip i ta te one o r t h e o ther .

A t t empera tu re s above about 300°c, increased pressure and added s a l t grea t ly increase t h e solubility of quartz. However, when a solution is hea t ed a t cons tant pressure, eventually i t will e i ther boil or en t e r a f ie ld of r e t rog rade qua r t z solubility. In e i ther event , qua r t z should prec ip i ta te , decreas ing t h e permeabili ty in t h e ho t t e s t par t of a convecting hydrothermal system. This precipitation of qua r t z is likely t o occur a t 300' t o 550°c, depending on t h e dep th of c i rcula t ion and t h e salinity of t h e system. An impermeable barr ier may form by a combination of

qua r t z deposit ion and quasi-plastic flow t h a t prevents subsequent d i r ec t c o n t a c t of circulating wa te r a t hydros ta t ic pressure with very hot rock, o r magma. This, in turn , will i nc rease t h e t ime required t o cool a shallow, in t rus ive magma, and might lead t o very s t e e p t empera tu re and pore-pressure gradients ac ros s t h e impermeable barrier. If t h e impermeable barr ier is then ruptured by se ismic ac t iv i ty or increasing pore pressure in t h e confined, high-pressure side of t h e system, a hydrothermal explosion may occur. This may be a mechanism by which some breccia pipes form. T h e sudden dec rease in density of t h e pore fluids, format ion of s t eam, separa t ion of gases, and decrease in t empera tu re accompanying t h e hydrothermal explosion should cause si l ica and K-rich feldspar t o prec ip i ta te a long with o ther minerals.

ACKNOWLEDGMENTS

Por t ions of th is manuscr ip t were reviewed by B. R. Berger , 3. W. Hedenquist , and D. E. White. I t has benef i ted grea t ly f rom their comment s and suggestions.

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APPENDIX

Information For Use In Calculating Silica Solubilities

Approximate solubilities of selected silica species in liquid water a t the vapor pressure of the solution can be calculated using equations (a)-(f) below (after Fournier, 1981; Fournier and Potter, 1982b). Concentrations of dissolved silica (S) are in rng/kg, t is temperature in degrees Celsius and the temperature b range of application is 0' to 250 C, except as noted.

Quar tz Log S = [-1309/(t+273.15)] + 5.19 ( a )

Chalcedony Log S = [-1032/(t+273.15)] + 4.69 ( b )

a - C r i s t o b a l i t e

Log S = [-1000/(t+273.15)] + 4.78 ( c )

6 - C r i s t o b a l i t e

Log S = [ -781/( t+273.15)] + 4.51 (d l

h r p h o u s s i l i c a

Log S = [ -731/( t+273.15)] + 4.52 ( e l

Q u a r t z (20'-330'~)

t = -42.196 + 0.28831 S - 3.6685 x s2

+ 3.1665 x s3 + 77.034 log S ( f

More precise solubilities of amorphous silica in the ternperature range 90' t o 3 4 0 ' ~ a t the vapor pressure of the solution and a t 1000 bars can be calculated using equations (g) and (h) (from Fournier and Marshall, 1983). T is temperature in Kelvin and m is the molality of dissolved silica.

Vapor Pressure of solution

log m = -6.1 16 + (0.01625 T) 41.758 x T2)

+ (5.257 x T3)

1000 bars

Appendix Table 3.Al--Temperatures, enthalpies (Keenan et al., 1969), and quartz solubilities (Fournier and Potter, 1982b) in liquid and gaseous water (steam) at the vapor pressure of the solution

No. T H Si02 No. T H Si02 No. T H Si02

OC J /g mg/kg OC J /g %/kg OC J/g mg/kg

log m = -7.010 + (0.02285 T) 43.262 x T2)

The solubility of quar tz in water in t h e t empera tu re range 25' t o 900°C a t specific volume (V) of t he solvent ranging from about I t o 10 and from 300° t o 6 0 0 ' ~ at specific volume of t h e solvent ranging f rom about 10 to 100 can be calculated using equations (i)-(I) (from Fournier and Pot ter , 1982a).

where

The solubility of quar tz in saline solutions can be calcula ted using equation (i) when (-log p F) is subst i tu ted fo r (log V), where P is t he density of the solution and F is t h e weight f rac t ion of water in t h a t solution (Fournier, 1983b).

The molal solubility of amorphous sil ica in saline solutions (m,) a t temperatures ranging f rom 100° t o 340°c, and pressures ranging from the vapor pressure of t h e solution t o about 1000 bars can be calcula ted using the following equations where ps is t h e density of t h e saline solution,pO is t h e density of pure wa te r a t any given temperature and t h e indicated pressure (obtained f rom s team tables) and m0 is t h e molal solubility in pure water obtained e i ther from equation (g) for t h e vapor pressure of t h e solution, or equation (h) for 1000 bars (Fournier and Marshall, 1983).

l o g m , = - n log psF

+ n l o g pO(v.p . ) + l o g m0 (v.p.1 (m)

Appendix Table 3.A gives t empera tu re , enthalpy, and quar tz solubility d a t a t h a t a r e useful for calculating silica concentra t ions a f t e r mixing of wa te r s with d i f ferent init ial t empera tu res and a f t e r boiling, a s discussed by Fournier and P o t t e r (1982b).

The dissociation of silicic acid a s a function of pH can be determined a s follows, where square brackets denote ac t iv i t ies of t h e indicated species, K I is t he f i rs t dissociation constant , m is molality,Y is t h e ac t iv i ty coefficient, and T is t empera tu re in Kelvin

Values of YHjSi0; can be ca lcula ted using the Debye Hiickel equation. When mH4s i04 is determined using equation (i) and substi tuting -log p F fo r log V, the value of YHQSi is unity.

Values o? K in t h e t empera tu re range 1 t o 350°C can be calcuia ted using t h e following equation

+ .000133266 T~ + 267.6478 l o g T ( r )

For most natura l thermal wa te r s values of pH a r e too low for t h e second dissociation of si l icic acid t o be important, and

n = ( n )

[ l o g pO(lOOO b a r ) - l og P O ( V . ~ . ) ]

Chapter 4 CARBONATE TRANSPORT AND DEPOSITION

IN THE EPITHERMAL ENVIRONMENT Robert 0. Fournier

INTRODUCTION

The factors af fect ing the transport and deposition of carbonate in hydrothermal sys tems have been discussed in detail by Holland and Malinin (1979). Solubilities of carbonates a r e strongly influenced by pH, PCOIl temperature , and the presence of o ther disso ved salts. The alkali carbonates, Na, K, and Li, a r e relatively soluble a t a l l temperatures and generally precipi ta te only where the re is ex t r eme evaporation. In contras t , t he alkaline e a r t h carbonates, Ca , Mg, Sr, and Ba, a re moderately t o sparingly soluble and commonly precipi ta te in hydrothermal systems. Ca lc i t e is by f a r t h e most abundant and important carbonate found in the epi thermal environment, and more solubility da ta a t hydrothermal conditions a r e available for i t t han for any of t h e o ther carbonates. Therefore, a f t e r briefly reviewing the system C02-water , t he discussion will focus on t h e t ranspor t and deposition of ca lc i te in hydrothermal solutions. The behaviors of o ther moderately t o sparingly soluble carbonates in hydrothermal solutions a r e similar t o t h a t of calcite.

given in Table 4.3. Adding sa l t t o the system C02- H 2 0 increases t h e Henry's Law coefficient and decreases the solubility of carbon dioxide in the solution (Fig. 4.1). Salting-out coefficients for carbon dioxide in sodium chloride solutions generally a r e of t he Steschenow type

where k is t he salting-out coefficient, m is t he molali ty of NaCl, and KOH and KH a r e respectively t h e Henry's Law coeff ic ients for pure water a s solvent and fo r t h e saline solution (Ellis and Golding, 1963). Approximate salting-out coefficients can be obtained from equation (b) in Table 4.1, and a t se lec ted t empera tu res f rom Table 4.3.

Some dissolved carbon dioxide r eac t s with water (hydrates) t o form carbonic acid

and some of t h e carbonic acid, in turn, dissociates according t o t h e react ions

C 0 2 DISSOLVED IN AQUEOUS SOLUTIONS H p 3 = HDj + Hf ( 4 )

There is extensive l i t e r a tu re on pressure-volume- t empera tu re measurements for t h e system C02-water , with and without additional dissolved sa l ts (Bowers and Helgeson, 1983; and references therein). The experimental work t h a t is most applicable to conditions appropr ia te for t h e format ion of epi thermal o re deposits was carr ied ou t by Ellis and Golding (1963), who used solubility d a t a t o ca lcula te Henry's Law coefficients, KH, fo r carbon dioxide in water and NaCl solutions (Fig. 4.1). According t o Henry's Law

where fCO2 is t h e fugaci ty of carbon dioxide and X is t h e mole f rac t ion of carbon dioxide dissolved in t h e liquid phase. Because fugacity coefficients for carbon dioxide in dilute aqueous solutions a t temperatures below about 3 3 0 ' ~ a r e near unity (Ellis and Golding, 19631, fC02 in equation ( I ) can be replaced by the partial pressure of carbon dioxide, PEOZ, with l i t t l e error. An equation expressing Henry's aw coeff ic ient for t h e system carbon dioxide-water a s a function of temperature is given in Table 4.1, equation (a). The constants used in t h e equations of Table 4.1 a r e given in Table 4.2. Values of t he Henry's Law coefficient for t h e system C 0 2 - H 2 0 a t s e l ec t ed temperatures a r e

0 0 100 200 300 400

Temperature , OC

Figure 4.1. Values of Henry's Law constant, %, for the solution of carbon dioxide in water and sodium chloride solutions. (Redrawn from Ellis and Golding, 1963).

Tab le 4.1--Equations e x p r e s s i n g t h e t empera tu re dependence of v a r i o u s e q u i l i b r i u m c o n s t a n t s and o t h e r c o e f f i c i e n t s t h a t a r e d e s c r i b e d i n t h e t e x t . T i s t e m p e r a t u r e i n k e l v i n o r OK, and t i s t empera tu re i n Ce l s ius o r OC. The r e s p e c t i v e c o n s t a n t s f o r e q u a t i o n s ( a ) t h rough (k ) a r e g iven i n Table 4.2. The d a t a used i n t h e d e r i v a t i o n of t h e s e e q u a t i o n s a r e shown i n Tab le 4.3.

KH = a + bT + CT' + dT-I + e log T ( a )

k = a + bT + cT2 + d~~ + e l o g T ( b )

- l o g K~ = a + b~ + CT-I + d l o g T ( c )

- l o g K~ = a + b~ + c ~ ' + d ~ - l + e l o g T ( d l

where square brackets indicate ac t iv i t ies of t he enclosed species, and KI, and K 2 a r e respectively the f i rs t and second dissocia t~on constants of carbonic acid (Table 4.3). Equations (c) and (dl in Table 4.1 express the t empera tu re dependence of K1 and K2. The reaction shown by equation (3) takes p lace relatively slowly, while the react ion shown by equation (4) is a lmost instantaneous (Kern, 1960). This information will be of use la ter when t h e consequences of boiling a r e discussed.

By tradit ion a distinction is not made between aqueous C 0 2 and H 2 C 0 3 , and to t a l dissolved C 0 2 is repor ted a s H2C03. A ne t react ion is generally wri t ten

- l o g K~ = a + bT + cT2 + d ~ - ' + e l o g T (el 032(gas) + = H2033 ( 8

HL = a + b t + c t 2 + d t 3 + e t 4

+ f t 5 + gt-l + ht-' + i l o g t ( f )

HG = a + b t + c t 2 + d t 3 + e t 4

+ f t 5 + gt-l + ht-' + i l o g t ( g )

- l o g Kc = a + bT + cT2 + dT3 + eT-l (h )

A = a + b~ + C T ~ + d ~ '

+ e~~ + fT-' + g l o g T ( i )

- l o g = a + b~ + c ~ '

+ dT3 + eT-l + fT-' (k )

where KO is t he equilibrium constant for t h e react ion shown by equation (8). Values of KO a t t empera tu res ranging f rom loo0 t o 3 0 0 ' ~ (Table 4.3) can be calculated using equation (e) in Table 4.1. The [H2C03] t e r m in equations (5) and (9) includes the ac t iv i ty of dissolved, nonhydrated C 0 2 .

H 2 C 0 3 is less ionized a t high t empera tu res compared t o I w temperatures (value K1 range f rom about 10'g*57 a t OOC to about 10-8.44at 3 0 0 ~ ~ ) . Therefore, a s a hydrothermal solution cools, bicarbonate dissociates (equation (4)), l iberating hydrogen ions t h a t a t t a c k t h e minerals in t h e wall rock. Hydrolysis react ions involving feldspars generally buffer t h e pH a t near neutra l t o slightly ac idic conditions and cooling solutions become richer in cations (Na a t higher t empera tu res and C a a t lower temperatures) a s hydrogen ions a r e consumed by the formation of micas or clays.

Where boiling occurs, generally a s a result of decreasing hydrostatic pressure exe r t ed upon an ascending hydrothermal solution, C 0 2 is strongly parti t ioned in to t h e gas (steam-rich) phase. The to t a l gas pressure is equal t o the sum of t h e par t ia l pressures of all t h e const i tuent gases, and these partial pressures a r e proportional t o the respect ive mole fractions. By Raoult's Law, for a mixture of two components t h a t exhibit ideal behavior (may b e closely approximated when t h e components a r e similar in rnolecular s t ructure , or when one of the components i s present in g r e a t excess)

Table 4.2--Coefficients for use with equations listed in Table 4.1

- log K1

- log K2

- log KO

HL

HG

- log Kc

A

B

- log AO

K~

k

- log K1

- log K2 - log KO

H~

HG

- log Kc

A

B

- log AO

where PYa and Pxb a r e the vapor pressures of pure in the liquid exsolves quickly in to t h e gas phase, while components A and B respectively a t t h e given only a small amount of water changes t o s team. With temperature , and Xa and X b a r e t h e respect ive mole continued boiling, t he mole f rac t ion of C 0 2 in t h e gas f rac t ions of A and B ~n t h e m ~ x t u r e . phase steadily decreases because l i t t l e a d d ~ t ~ o n a l C 0 2

When boiling is f i r s t init iated, t h e r a t io of C 0 2 is available t o parti t ion in to t h e gas phase while t h e t o wa te r in t h e gas phase tends t o b e relatively l a rge f rac t ion of water t h a t is conver ted t o s t e a m increases because most of t h e carbon dioxide initially dissolved a t a relatively constant ra te . As t h e t empera tu re of

Table 4.3--Values of d i s s o c i a t i o n c o n s t a n t s , e n t h a l p i e s of l i q u i d w a t e r HL and s team . . HG, and Debye-Huckel C o e f f i c i e n t s , A and B , f o r t h e i n d i c a t e d t empera tu res .

En tha lpy u n i t s a r e J /g . Un i t s f o r A a r e kg1/2mole-1/2 and f o r B a r e

kg1/2mole-1crn-1 x l o 8 ( t h e product x B c a n c e l s t h e l o 8 f a c t o r )

~ ( O C ) 100 125 150 175 200 225 250 275 300 Reference

* E x t r a p o l a t e d . ( 1 ) Henley e t a l . (1984) . (2 ) E l l i s and Golding (1963) . ( 3 ) Keenan e t a l . (1969).

t he ascending gas-water mixture decreases, t h e volume of t h e gas phase increases due t o t h e decrease in hydrostatic load. The n e t e f f e c t is a dras t ic decrease in the partial pressure of C 0 2 a s a boiling fluid ascends toward the earth 's surface.

Procedures for calculating t h e parti t ioning of relatively volati le consti tuents between coexisting liquids and gases, using hand-held, programmable calculators, a r e described in Henley e t al. (1984). For relatively dilute sys tems and low initial dissolved gas concentrations, a distribution coefficient, B, is defined a s the concentration of gas in t h e vapor divided by t h e concentration of gas in t h e liquid. Ciggenbach (1980) derived the following equation t h a t expresses the t empera tu re dependence of B for carbon dioxide in dilute aqueous solutions

l o g B = 4.7593 - .01092t ( 1 1 )

decompression (single-step s t eam separation)

where Co is t h e init ial concentration of dissolved C 0 2 before boiling, B is t h e distribution coefficient, and y is t h e f rac t ion of separa ted steam. The corresponding equation t h a t gives t h e concentration of C 0 2 in t h e coexisting s t eam ICv) is

Values of y a r e generally ca lcula ted using enthalpy d a t a for pure boiling water and t h e relationship

where t is t empera tu re in d e rees Celsius. Equation 8 (11) is valid f rom 100' t o 340 C. Henley e t al. (1984) Ho- HL

Y = (14 ) give t h e following equation for calculating t h e concentration of C 0 2 remaining in the liquid phase

Hc - % (CI) for t h e situation in which all t h e evolved gas where Ho is t h e enthalpy of the initial liquid prior t o remalns in con tac t with a boiling fluid during adiabat ic boiling, and HL and HG a r e t h e enthalpies of

THE SOLUBILITY OF CALCITE IN AQUEOUS SOLUTIONS

coexisting liquid water and s team a f t e r boiling o (Table 4.3). Enthalpies of liquid water and s t eam a r e generally obtained from s t eam tables (Keenan e t al., 1969) or they can be calcula ted using equations (f) and (g) in Table 4.1. Equation (131, however, yields values of y t h a t a r e slightly in er ror because t h e enthalpy of s t eam containing C 0 2 is d i f ferent from the enthalpy of pure s team. Other f ac to r s also may cause the ca lcula ted concentra t ion of C 0 2 in the liquid and s team fract ions of a boiling solution t o be in error. -1

Assulnptions implicit in t h e use of equations (12) and (13) a r e t h a t dissolved C 0 2 does not become supersa tura ted in the liquid phase a s pressure is 0" 1 released, and t h a t l i t t le or no H C 0 3 conver ts t o 0

H C O j a s the boiling solution cools. $he rapid t ransfer a of most of t h e dissolved C 0 2 in to the s t eam fract ion - a t an ear ly s t age of boiling and the relatively slow conversion of dissolved C 0 2 to H2C03 (previously discussed) will tend t o l imi t t he amount of H C O t h a t -2

can form, but some non-equilibrium parti t ioning of C 0 2 between the liquid and gas phase is likely, particularly when the f i rs t boiling is in i t ia ted a t a ternperature below about 2 0 0 ' ~ . Another f ac to r t h a t must be considered is physical removal of t h e s team fract ion from con tac t with the residual liquid a s the boiling process proceeds. Compared t o single-step s t eam separation, multistep and continuous s t eam

Ellis (1959, 1963) determined experimentally the solubility of ca l c i t e in aqueous solutions a t conditions appropriate for t he format ion of epi thermal ore deposits. A t a given partial pressure of C 0 2 , t he solubility of ca lc i te decreases with increasing t empera tu re (Fig. 4.3). Adding NaCl t o t h e solution increases the solubility of ca lc i te (Fig. 4.4). A t any given t empera tu re t h e solubility of ca lc i te in solutions in equilibrium with a vapor phase increases with increasing C O pressure until mc-2 = 1 molelkg (Miller, 1952; ?egnit e t al., 1962). In solutions held a t a constant t o t a l pressure, t h e solubility increases with increasing C 0 2 concentra t ion until rnc02 21 mole/kg and then decreases a t higher C 0 2 concentrations (Sharp and Kennedy, 1965; Malinin and Kanukov, 1971).

The simplest equation representing the react ion by which ca l c i t e dissolves in aqueous solutions can be wri t ten

-

-

and t h e equilibrium constant (Kc) for reaction (15) is

separation in contact. t he las t Henley result liquid in e t and much al. s t eam (1984) lower fractions concentrations present methods t h a t of a r e C 0 2 and in -3 . l o o && I Temperature 200 I I OC 300 I I 400 -

equations for dealing with multistep and continuous s team separation. Figure 4.2 shows values of C,/Co Figure 4.2. The C02 remaining in the for single-step s t eam separa t ion for a var ie ty of inltlal residual liquid (Cl) after single-step and final temperatures , ca lcula ted using equations (12) steam separation at various temperatures to and (14). C02 ,in the initial liquid (Co) before

boiling.

Equation (16) is useful mainly for tes t ing whether a solution of given cornposition is unsaturated, sa tura ted, or supersa tura ted in respect t o calcite. Values of Kc in t h e t empera tu re range 100' t o 300°C (Table 4.3) can be calculated using equation (h), Table 4.1. Because very l i t t l e CO? is present in most natura l hydrothermal solutions, t h e solubility of ca l c i t e is commonly expressed in t e rms of react ions involving H', HCO, and f using equations (5), (71, (9), and (16). C02

- 100 150 200 250 300

Temperature. OC E .- - - .-

F i g u r e 4.3. The s o l u b i l i t y o f calcite i n w a t e r = 2

u p to 300% a t v a r i o u s partial p r e s s u r e s o f c a r b o n d iox ide . (Redrawn f rom E l l i s , 1959).

n

In equations (16), (181, (20), and (221, t h e ac t iv i ty of ca l c i t e is unity if t he re i s no significant substi tution of o ther cations fo r calcium in solid solution, such a s Mg, Fe , or Mn.

In order t o evaluate equations (16)-(221, ac t iv i t ies of t h e indicated aqueous species must be used. In dilute solutions, ac t iv i t ies of dissolved const i tuents a r e about equal t o t h e corresponding molalities. In saline solutions, however, t he molali ty of e a c h species i (mi) must be multiplied by i t s ac t iv i ty coeff ic ient (Y.) t o obtain the ac t iv i ty (ai = Y im.). Activity coeff ic ients for solutions with ionic s t rengths less than about 2 molal can be ca lcula ted using an extended form of t h e Debye-Hiickel equation

where gi is t h e ionic charge, I t he ionic s t rength , and A, B, ai and b a r e constants (Henley e t al., 1984). However, A and B vary with temperature . Their values from 100' t o 350°c, in 2 5 ' ~ increments, a r e given in Table 4.3. The coefficients A and B a lso can be calcula ted t o t h r e e decimal places using equations (i) and (j), Table 4.1. The ionic s t rength is defined a s

For most hydrothermal waters I is approximately equal t o t h e sum of m~ + and m . Values of Si and zi a r e l isted in Table 4.C Up to%>O°C, b has values in t h e range 0.03 t o 0.05 when concentrations a r e up t o 3 molal (Helgeson, 1969).

In natura l hydrothermal solutions many dissolved const i tuents and a variety of chemical react ions involving solids, liquids and gases influence t h e

Temperature, OC

F i g u r e 44. The s o l u b i l i t y o f calcite i n w a t e r a n d s o d i u m c h l o r i d e s o l u t i o n s a t a c a r b o n d i o x i d e p r e s s u r e o f 1 2 a t m o s p h e r e s (12.2 b a r s ) . (Redrawn f r o m E l l i s , 1963).

dissolution and deposition of calcite. The situation is particularly complex when boiling occurs. Computer programs can be used t o eva lua te these complex react ions (Truesdell and Singers, 1971, 1974; Morel and Morgan, 1972; Kharaka and Barnes, 1973; Truesdell and Jones, 1974; Plummer e t al., 1975; Wolery, 1979; Reed, 1982; Reed and Spycher, 1984). These programs, however, generally require main-frame computers for thei r execution.

Arnorsson (1978) specifically ca lcula ted the amount and location of ca l c i t e deposition in geothermal wells in Iceland where natura l thermal waters flash t o s t eam during production of t h e resource. The results of his calculations, showing the degree of supersaturation with respect t o ca l c i t e t h a t occurs during single-step adiabat ic flashing (boiling) in the wells, a r e shown in Figure 4.5 (supersaturated solutions t h a t plot below t h e thick solid line).

Programs t h a t can be used with programmable hand-held calculators, and t h a t a r e applicable to ca lc i te transport and deposition in natura l waters, a r e given in Henley e t al. (1984). These programs a r e very useful even though a few simplifying assumptions a r e required, and only t h e most important dissolved species a r e included. The ensuing discussion follows t h e general procedures presented in Henley e t al. (1984).

From a consideration of t h e cation-anion charge balance t h a t must be maintained in a l l solutions (and neglecting e f f ec t s of oxidation-reduction react ions and pH-dependent ions t h a t a r e likely t o be present only in very small amounts in most natura l hydrothermal solutions, such a s CO;, H2SiOi, and OH-) a constant,A,

0 Table 4.4--Values of i o n i c cha rge , z , and i o n s i z e pa ramete r , a , f o r t h e

common i o n i c s p e c i e s i n geothermal f l u i d s (from Henley e t a l . , 1984)

H+ ~ a + HCO; HS- H3Sio; H ~ B O ; F- SO; N H ~ HSO;

OH- CO; a- ~ i + K+ ~ a + + M ~ + +

150 200 250 150 200 250

T e m p e r a t u r e O C

F i g u r e 4.5. T h e -computed a c t i v i t y p r o d u c t o f ~ a + + a n d Cog i n g e o t h e r m a l w a t e r s d u r i n g s i n g l e - s t e p adiabatic f l a s h i n g i n r e l a t i o n t o the calc i te s o l u b i l i t y c u r v e ( t h i c k solid l i n e ) . The s o l i d l i n e s assume maximum d e g a s s i n g a n d t h e d a s h e d l i n e s 1 / 5 o f maximum degass ing. (From Arnorsson, 1978).

can be defined t h a t i s independent of t empera tu re and equal t o t h e sum of t h e principal pH-dependent ions

The concentrations of t h e ionic species indicated in equation (25) can be calcula ted using t h e following relationships, where K 1 is t h e f i r s t dissociation constant for each of the respect ive weak acids

%2( t o t a l ) "k"03 = (26 )

[H+l Y m j

The reader is referred t o Henley e t al. (1984) for derivation of equations (261429).

If t he pH, ionic s t rength , and chemical composition (particularly to t a l dissolved carbon, silica, boron, and ammonia) of a solution a r e known a t a given temperature , equations (25)-(29) can be used t o

es t ima te the indicated species concentrations a t any other t empera tu re up t o t h e l imit of t he available thermodynamic d a t a (now about 300' t o 350°C). Note, however, t h a t equations (25)-(29) do not t a k e account of changing concentra t ions and parti t ioning of const i tuents between t h e liquid and gas phase during boiling.

For adiabat ic boiling resulting from decompression, t h e value of 'changes a s a function of the f rac t ion of s t eam (y) t h a t forms.

' ( b e f o r e b o i l i n g ) = - Y ) ' ( a f t e r b o i l i n g )

(30)

The e f f e c t of parti t ioning of C 0 2 between liquid and gas can be accounted fo r by using the relationship

where n and nv a r e t h e number of moles of t he indicated species in t h e liquid and gas phases respectively, and values of AQ a t various ternperatures with boiling water and bo~ l ing NaCl solutions a s solvents a r e given in Ellis and Golding (1963). A, fo r dilute solutions also can be calculated using equatlon (k) in Table 4.1. Rearranging equation (31) and substi tuting

For one-step s t eam separation without carbonate precipitation (supersaturation allowed in the calculation) the to t a l number of moles of C02-bearing species remains constant during the boiling process, even though C 0 2 parti t ions between the gas and liquid phases

The concentration of CO; in equation (33) cannot be neglected when solutions boil because t h e pH may increase significantly (Fig. 4.6). Combining equations (32) and (33)

and dividing equation (34) by nHCOj

Equation (35) is expressed in t e rms of mole ratios, so molal concentra t ion units can be substi tuted for t he number of moles of t he given species in the liquid fraction. Combining equations (51, (71, and (351, multiplying molali t ies by act iv i ty coefficients t o obtain ac t iv i t ies where required, and rearranging gives

Account can be taken of the parti t ioning of o ther volati le const i tuents in a similar manner. The resulting equations can be used to e s t ima te by

Hveragerdi

Namafall

Nesjavellir p I Leira

Reyk janes

Figure 4.6. The variation in pH in geothermal waters during one-step adiabatic flashing in relation to the calcite solubility curve (thick solld line). The solid lines assume maximum degassing and the dashed lines 1/5 of maximum degassing. (Redrawn from Arnorsson, 1978).

R. 0. FOURNIER 7 1

i tera t ion t h e pH and distribution of pH-dependent species a t given t empera tu res for boiling conditions. The equations and procedures a r e cumbersome using a hand-held ca lcula tor , but a r e easily deal t with using table-top micro o r personal computers.

Where cooling occurs adiabatically (generally too quickly fo r solution-mineral reactions involving sil icates t o buffer pH) t h e pH of t h e residual liquid usually rises a s a resul t of parti t ioning of C O and other acid-forming gases in to t h e s team phase. Agure 4.6 shows calcula ted changes in pH tha t accompany single-step adiabat ic flashing of natura l thermal waters in geothermal wells in Iceland (Arnorsson, 1978). These a r e t h e same well waters used to i l lustrate the degree of supersaturation with respect t o ca lc i te t h a t occur s during single-step flashing (Fig. 4.5).

SUMMARY

In most natura l waters heating will cause ca lc i te and other carbonates t o precipitate, whereas cooling without boiling will cause them t o dissolve (Fig. 4.4). However, where a n ascending solution boils a s a result of decompression (cooling adiabatically), carbonate is likely to precipi ta te a s a result of t he boiling (Fig. 4.5). The cooling t h a t tends t o move a solution toward a condition of undersaturation with respect t o the various carbonate minerals is generally more than offse t by the s t rong parti t ioning of C 0 2 in to the vapor phase (and concomitant dec rease in partial pressure of C 0 2 ) t h a t decreases the solubility of carbonates. A t present, calculations t h a t t a k e account of all t he physical processes (mainly boiling) and chemical react ions t h a t influence t h e transport and deposition of carbonate minerals can be carr ied ou t only with the aid of large computers. However, if only t h e most abundant dissolved species in natura l waters a r e considered, and simplifying assumptions a r e made about enthalpies of coexisting liquids and gases (mainly steam), small table-top cornputers and hand-held, programmable ca lcula tors can be used effect ively t o ca lcula te t h e approximate conditions fo r transport and deposition of ca l c i t e in hydrothermal solutions.

REFERENCES

Arnorsson, S., 1978, Precipitation of ca lc i te from flashed geothermal waters in Iceland: Contributions t o Mineralogy and Petrology, v. 66, p. 21-28.

Bowers, T. S., and Helgeson, H. C., 1983, Calculation of t h e thermodynamic and geochemical consequences of nonideal mixing in the system H20-CO NaCl on phase relations in geologic sys t ems :2 -~qua t ion of s t a t e fo r H20-C02-NaCI fluids a t high pressures and temperatures: Geochimica et Cosmochimica Acta , v. 47, p. 1247-1275.

Ellis, A. J., 1959, The solubility of ca l c i t e in carbon dioxide solutions: American Journal of Science,

Ellis, A. J., 1963, The solubility of ca lc i te in sodium chloride solutions at high temperatures: American Journal of Science, v. 261, p. 259-267.

Ellis, A. J., and Golding, R. M., 1963, The solubility of carbon dioxide above 1 0 0 ~ ~ in water and in sodium chloride solutions: American Journal of Science, v. 261, p. 47-60.

Giggenbach, W. F., 1980, Geothermal gas equilibria: Geochimica e t Cosmochimica Acta , v. 44, p. 2021-2032.

Henley, R. W., Truesdell, A. H., and Barton, P. B., Jr., 1984, Fluid-mineral equilibria in hydrothermal systems: Society of Economic Geologists, Reviews in Economic Geology, Volume 1, 267 p.

Holland, H. D., and Malinin, S. D., 1979, On t h e solubility and occurrence of non-ore minerals; & Barnes, H. L. (ed.), Geochemistry of Hydrothermal Ore Deposits (2d ed.): John Wiley and Sons, New York, p. 461-508.

Keenan, J . H., Keyes, F. G., Hill, P.G., and Moore, 3. G., 1969, S team Tables (international edition- m e t r i c units): John Wiley and Sons, New York, 162 p.

Kern, D. M., 1960, The hydration of carbon dioxide: Journal of Chemist ry Education, v. 37, p. 14-23.

Kharaka, Y. K., and Barnes, I., 1973, SOLMNEQ: Solution-mineral equilibrium computations: National Technical Information System Technical Repor t PB 214-899, 82 p.

Malinin, S. D., and Kanukov, A. B., 1971, The solubility of ca l c i t e in homogeneous H20-NaCI-C02 sys tems in the 2 0 0 ~ - 6 0 0 ~ ~ t empera tu re interval: Geochemistry International, v. 9, p. 410-418.

Miller, J. P., 1952, A portion of the system calcium carbonate-carbon dioxide-water, with geologic implications: American Journal of Science, v. 250, p. 161-203.

Morel, F., and Morgan, J., 1972, A numerical method for computing equilibria in aqueous systems: Environmental Science and Technology, v. 6, p. 58-67.

Plummer, L. N., Parkhurst , D. L., and Kosiur, D., 1975, MIX2: A computer program for modeling chemical reactions in natural waters: U.S. Geological Survey, Water-Resources Investigations, p. 75-61.

Reed, M. H., 1982, Calculation of multicomponent chemical equilibria and reaction processes in sys tems involving minerals, gases and a n aqueous phase: Geochimica e t Cosmochimica Acta , v. 46, p. 513-528.

Reed, M. H., and Spycher, N., 1984, Calculation of pH and mineral equilibria in hydrothermal waters with application t o geothermometry and studies of boiling and dilution: Geochimica e t Cosmochimica Acta, v. 48, p. 1479-1492.

Segnit, E. R., Holland, H. D., and Biscardi, C. J., 1962, The solubility of ca l c i t e in aqueous solutions-I. The solubility of ca lc i te in water between 75' and 200' a t COC? pressures up t o 60 a tm: Geochimica e t osmochimica Acta, v. 26, p. 1301-1331.

Sharp, W. E., and Kennedy, G. C., 1965, The sys tem C a 0 - C 0 2 - H 2 0 in t h e two-phase region ca l c i t e and aqueous solution: Journal of Geology, v. 73, p. 391-403.

Truesdell , A. H., and Jones, B. F., 1974, WATEQ, a compu te r program fo r ca lcula t ing chemical equil ibria of na tura l waters: U.S. Geological Survey, Journal of Research, v. 2, p. 233-248.

Truesdell , A. H., and Singers, W. A., 1971, Compute r ca lcula t ion of downhole chemis t ry in geothermal

areas: New Zealand DSIR Chemis t ry Division Repor t CD2136, 145 p.

Truesdell , A. H., and Singers, W. A., 1974, Calcula t ion of aqui fer chemis t ry in hot-water geothermal systems: U.S. Geological Survey, Journal of Research, v. 2, p; 271-278.

Wolery, A. T., 1979, Calcula t ion of chemical equilibrium between aqueous solutions and minerals; t h e EQ3/6 so f tware package: UCRL-52658, Lawrence Livermore Laboratory.

Chapter 5 FLUID-INCLUSION SYSTEMATICS IN EPITHERMAL SYSTEMS

R. J. Bodnar, T. J. Reynolds, and C. A. Kuehn

INTRODUCTION

Fluid-inclusion analyses have provided some of the most useful information for determining the physical and chemical environments of mineral formation. The purpose of this chap te r is t o describe those fluid-inclusion character is t ics which se rve t o distinguish relatively near-surface, epi thermal formation conditions f rom deeper and, potentially, higher temperature formation conditions, and t o discuss several techniques and problems which a r e specific t o fluid inclusions trapped in t h e epi thermal environment. A detailed summary and cr i t ique of fluid-inclusion l i tera ture re la ted t o epi thermal sys tems has not been a t tempted. For th is information t h e reader is referred t o the recent compilations of Buchanan (19811, Heald-Wetlaufer et al. (1983), Roedder (19841, and Hedenquist and Henley (19 85). Moreover, we have not a t t e m p t e d t o r e l a t e any particular fluid-inclusion cha rac te r i s t i c t o a specific type o r s t age of mineralization, because a n adequate d a t a base t o do so does not presently exist.

This presentation is l imited t o t w o subjects--the petrography and petrology of fluid inclusions from t h e epithermal environment--and is in tended t o provide t h e explorationist with a basic understanding of t h e cr i ter ia for recognizing and in terpre t ing inclusions trapped in this environment. Two impor t an t topics will be discussed in detail: (1) t h e identification and in terpre ta t ion of fluid inclusions trapped f rom boiling fluids, and (2) t he identification of gases (mainly C 0 2 ) in fluid inclusions and t h e e f f e c t of volati les on calculated pressures and depths of trapping. We will not, however, discuss t h e important chemical consequences of boiling and dissolved volati les, a s these subjects a r e covered in deta i l in o the r chap te r s in this volume (see Henley, 1985, this volume; Henley and Brown, 1985, this volume; Fournier, 1985, this volume; and Reed and Spycher, 1985, this volume).

INFORMATION AVAILABLE FROM FLUID-INCLUSION PETROGRAPHY

Character is t ics of fluid inclusions t rapped in the epi thermal environment con t ra s t markedly with those of inclusions formed in deeper environments. This sect ion considers those f ea tu res which a r e diagnostic of shallow crus ta l environments and which a r e readily observable by anyone with access to a s tandard petrographic microscope. In addition, owing t o the na tu re of fluid inclusions t rapped in t h e epi thermal environment, particular ca re in t h e se lec t ion of fluid inclusions for detailed microthermometr ic analysis

must be practiced, and such precautions will be discussed. Diagnostic information common t o all minerals containing fluid inclusions formed in the epi thermal environment will be presented first , followed by a deta i led discussion of information available from petrographic observations of fluid inclusions in quartz.

Fluid inclusions in t h e epitherrnal environment typically contain only t w o phases a t room temperature--a low-salinity H 2 0 liquid phase and a vapor bubble. Daughter minera ls of hal i te and sylvite a r e notably absent in the epi thermal environment. Readily observable evidence for gases is iacking also: gases rarely occur a s condensed phases in fluid inclusions, and evidence of gases i s normally not found by sirnple crushing tes ts , fo r reasons described below. However, low concentra t ions of gases have been identified by capaci tance manometer and mass spect rometr ic techniques (Sommer e t al., 1985; Hedenquist and Henley, 1985).

Exceptions t o these general i t ies may occur when an epi thermal sys tem overprints an earlier, higher t empera tu re system, or v ice versa. For example, a t Summitville, Colorado, which is a ra ther high-level, fossil hydrothermal system within a volcanic dome (Perkins and Nieman, 1982), some early quar tz rarely contains a f ew isolated healed microfractures defined by vapor-rich H 2 0 + C 0 2 (270 mole % C O ) inclusions and/or healed microfractures defined by hajite-bearing inclusions with smal l vapor bubbles. These inclusions presumably contain samples of ear ly magmat ic fluids trapped before the development of the near-surface epi thermal system a t th is locality. Similarly, high- t empera tu re ( 360°C), vapor-rich H 2 0 + C 0 2 (270 mole % C 0 2 ) inclusions found in some act ive geothermal sys tems (e.g., in ce r t a in deep portions of t h e Geysers geothermal field) could be magmat ic fluids. Also, t he high concentra t ions of C 0 2 and CHq present in fluid inclusions f rom sediment-hosted gold deposits may not be temporally re la ted to the hydrothermal sys tem at tending gold mineralization. Certainly, some epi thermal sys tems must have been subsequently buried and subjected t o fluids of deeper origins, and in these cases COZ-bearing and/or salt- s a tu ra t ed inclusions may postdate the epi thermal inclusions.

Fluid inclusions a r e t rapped in many minerals formed in t h e epi thermal environment; quartz, sphalerite, ca lc i te , and f luor i te have yielded useful t he rmomet r i c data. Of these minerals, quar tz usually provides the most f e r t i l e opportunities for collection of in terpre table fluid-inclusion data. Jus t a s t h e megascopic crustiform banding of quar tz is

c h a r a c t e r i s t i c o f v o i d filling in t h e epithermal environment , s o a r e t h e microscopic t ex tu res in quartz diagnostic. F u r t h e r m o r e , the fluid-inclusion textures in q u a r t z vary s y s t e m a t i c a l l y in a manner t h a t permits genera l t he rma l c o n d i t i o n s t o be predicted from f e a t u r e s observed u n d e r the microscope. These sys t ema t i c va r i a t ions i n epithermal-quartz fluid- inclusion t e x t u r e s m a y re f l ec t t he temperature dependence of t h e k inet ics of dissolution and reprecipi ta t ion o f q u a r t z during "maturationn* of individual inclusions. Thus , the following discussion of t h e microscopic t e x t u r e s diagnostic of t he epithermal environment a r e l i m i t e d to those found in quartz.

*The t e r m "matura t ion" is used t o describe t h e process of dissolution a n d r ep rec ip i t a t ion of a host mineral surrounding a t r a p p e d fluid. The init ial "immature" inclusion is genera l ly l a r g e and very irregularly shaped. With t ime , t h i s l a r g e inclusion will neck down t o form numerous s m a l l e r , more regularly shaped inclusions, with t h e m o s t ma tu re inclusion obtaining t h e negat ive c rys t a l s h a p e of t h e host mineral. (See Roedder, 1984; figs. 2-15.)

A t t h e ou t se t , i t should be noted t h a t the microscopic t ex tu res displayed by qua r t z discussed below a r e not unique t o t h e epi thermal environment. Similar t ex tu res m a y o c c u r in qua r t z from deeper, h igher- temperature hydrothermal systems. To mit igate possible ambigui t ies in the final in terpre ta t ion of t h e t e x t u r e s observed, many samples must be col lec ted and observed t o gain a broad perspect ive of a l l f e a t u r e s present, and these da ta must be combined wi th fundamental geologic knowledge. As is t r u e fo r most geologic studies, concIusions should n o t b e formulated f rom a single observation, and th is is t r u e fo r any study involving fluid inclusions.

From a fluid-inclusionistls viewpoint, quar tz formed in t h e epi thermal environment can be divided in to two groups--quartz t h a t contains fluid inclusions large enough t o study bl .5pm) and quar tz t h a t contains very f ew o r no inclusions large enough t o study with a s tandard petrographic microscope equipped for to t a l magnifications a t leas t a s high a s 480X. Figure 5.IA shows th ree di f ferent types of quar tz commonly found in t h e epi thermal environment t h a t contain few, if any, inclusions large enough t o study. The lower par t of Figure 5.1A shows finely crystalline, euhedral qua r t z crystals t h a t grew contemporaneously f rom many nucleation sites. This t ex tu re is of ten noted a s l a t e vug-fillings, but may also be a resul t of replacement (silicification) of original wallrock. The finely crystall ine quar tz seldom contains fluid inclusions large enough t o study, but on t h e r a re occasion when appropriate inclusions a r e found (see below), homogenization t empera tu res a r e typically <200°c and melting points a r e usually between 0 and -2Oc. Numerous voids, bounded by quar tz crys ta l faces, occur in this type of quartz. Rarely, these spaces may be sealed t o t r a p large fluid inclusions, thus recording either t h e l a t e s t fluids from

which t h e qua r t z grew, or fluids t h a t circulated a t s o m e l a t e r period, perhaps even mill ions of years after t h e c rys t a l s formed. Such l a rge inclus ions generally leak dur ing heat ing runs, indicat ing t h a t they are s t ruc tu ra l ly weak; thus, a t t en t ion t o th i s inherent weakness is necessary when attempting homogenization t e m p e r a t u r e determinations. Fu r the rmore , because t h e inclusions a r e weak, mechanical s t resses (including t h e preparation procedure) may permit l a t e fluids to e n t e r the inclusions. If mel t ing points of such angular, in tercrys ta l l ine inclusions a r e t o be determined, the d a t a should be in t e rp re t ed with care. More typically, t hese in tercrys ta l l ine spaces conta in only air, and appea r a s dark, angular cavi t ies a s shown in t h e early qua r t z in t h e lower portion of Figure 5.lA.

The finely banded, botryoidal q u a r t z in the c e n t e r of Figure 5.1A is chalcedonic quar tz . Fluid inclusions have not been observed t o d a t e in this type of quar tz , but i t is suspected t h a t such quartz generally fo rms a t low temperatures .

The coarse-grained, c lear , euhedral crys ta ls at t h e t o p of Figure 5.1A a r e also common in the epi thermal environment. Generally, inclusions in such c l ea r , euhedral qua r t z a r e rare , and if present, are very smal l ( 5 3 ~ m in longest dimension). Occasionally, however, c lear , euhedral qua r t z may have inclusions large enough t o observe a s shown in Figures 5.16 and 5.1C. Figure 5.IC is an enlargement of a growth zone in t h e c e n t e r of Figure 5.18, and shows t h a t t he growth zone in t h e qua r t z i s defined by smal l fluid inclusions ( thus of def in i te primary origin) of irregular shape and containing variable liquid t o vapor volumetr ic phase ratios. The presence of one-phase, liquid-filled inclusions in th is growth zone indicates t h a t healing of t h e inclusions had continued a t low t empera tu res (<lOoOc) a f t e r t he vapor phase had nucleated within t h e cavi ty upon cooling, and t h a t t h e inconsistent liquid t o vapor ra t ios exhibited in t h e growth zone a r e probably not a result of inhomogeneous en t r apmen t of liquid and vapor phases during boiling. Similar textura l relationships can be seen c lear ly in Figures 5.IE and 5.IF. In these figures, many vapor-rich inclusions coexist with liquid-filled inclusions in a coarse-grained euhedral quar tz growth zone. The presence of many one-phase, liquid-filled inclusions indicates t h a t t h e vapor-filled inclusions probably resulted f rom continued healing and "neckingq' of t he long, stringy, dendrit ic inclusion cavities a f t e r a vapor phase had nucleated within the inclusions in a way t h a t separa ted t h e vapor phase plus some liquid in to some inclusions and only the liquid phase into o the r inclusions. This same process is responsible for t h e variable liquid t o vapor volumetric ratios observed in healed microfractures crossing clear, euhedral qua r t z a s shown in Figure 5.1D. Again, what distinguishes continued healing of inclusions a t low t empera tu res a f t e r nucleation of a vapor phase from entrapment of inclusions during fluid boiling is the observation t h a t one-phase, liquid-filled inclusions coexist with inclusions of variable liquid-to-vapor ra t ios within the individual plane. Of course, boiling could have occurred during formation of these textures , but continued "necking" a f t e r nucleation of another phase would mask evidence for boiling.

Figure 5 . 1 A-I.

Figure 5.1 J-R.

Figure 5.1. Photomicrographs showing mineralogical and fluid-inclusion textures common in quartz formed in the epithermal environment.

A). Three types of quartz from the epithermal environment that rarely contain fluid inclusions large enough to study. In this example, finely crystalline euhedral quartz containing numerous dark, air-filled voids (bottom) is succeeded by banded, btryoidal quartz (center), and finally by large, clear euhedral quartz crystals (top). Bar scale equals 200 micrometers.

B). Zoned clear, euhedral quartz crystals (center) overgrown by banded chalcedonic quartz (top). Bar scale equals 200 micrometers.

C). Enlargement of central portion of Figure 5.1B showing numerous primary fluid inclusions with variable liquid-to-vapr ratios resulting from necking at low temperature after formation of a vapor phase in the inclusions. Sar scale equals 10 micrometers.

D). Plane of secondary inclusions exhibiting variable liquid-to-vapor ratios (including one-phase liquid inclusions) resulting from necking after the generation of a vapor phase at low temperatures. Bar scale equals 10 micrometers.

E). Clear, euhedral quartz crystal containing numerous primary fluid inclusions oriented perpendicular to growth surfaces. Bar scale equals 200 micrometers.

F). Enlargement of portion of Figure 5.1E showing numerous large, stringy primary inclusions with variable liquid-to-vapor ratios. One-phase liquid inclusions in this zone indicate necking at low temperature after a vapor phase had nucleated in the inclusions. Bar scale equals 10 micrometers.

G) . Plane of secondary inclusions with consistent liquid-to-vapor ratios. Bar scale equals 10 micrometers.

H). Clear, euhedral quartz crystals showing growth zones defined by dark bands of primary fluid inclusions. Bar scale equals 200 micrometers.

I). Enlargement of portion of Figure 5.1H showing irregularly shaped, prima? fluid inclusions, most with consistent liquid-to-vapor ratios, that homogenize at 230+10 C. Bar scale equals 10 micrometers.

J). Portion of a zoned quartz crystal showing primary fluid inclusions defining growth surfaces. Bar scale equals 200 micrometers.

K), L). Enlargements of portions of Figure 5.1J showing irregularly shaped, grimary inclusions, most with consistent liquid-to-vapor ratios, that homogenize at 250+10 C. Bar scale equals 10 micrometers.

M ) . Large euhedral quartz crystal showing numerous primary fluid inclusions defining a former crystal surface. Bar scale equals 200 micrometers.

N). Enlargement of growth zone shown in Figure 5.1M containing liquid-rich fluid inclusions with consistent liquid-to-vapor ratios and vapor-rich inclusions with more variable 1 iquid-to-vapor ratios.

0). Healed microfracture in quartz containing all vapor-rich inclusions. Bar scale equals 10 micrometers.

P). Euhedral quartz crystal with primary fluid inclusions defining former growth surfaces. Bar scale equals 200 micrometers.

Q). Enlargement of portion of Figure 5.1P showing regularly shaped primary inclusions that homogenize at 270+10°~. Bar scale equals 10 micrometers.

R). Original epithermal quartz that has been subjected to a regional metamorphic event. Although the original epithermal textures are still preserved, the quartz is cut by numerous microfractures defined by small C02-bearing inclusions introduced during metamorphism. Bar scale equals 200 micrometers.

Homogenization t e m p e r a t u r e s col lec ted f rom groups of pr imary inclusions o r planes of secondary inclusions exhibit ing var iable liquid-to-vapor volumetr ic phase ra t ios due t o "necking" a f t e r nucleation of t h e vapor phase will be erroneous, and therefore , misleading. Hence , i t is advisable t h a t homogenization t e m p e r a t u r e s not be col lec ted on such inclusions, al though mel t ing points of ice could be valid. The ca re fu l inclusionist should painstakingly survey samples t o l o c a t e t h e rare , isolated plane(s) or area(s) within a growth zone where t h e fluid inclusions exhibit consistent liquid-to-vapor volumetr ic phase ra t ios and only de te rmine homogenization tempera- t u re s f o r t hese inclusions. F o r th is t ype of qua r t z t h a t conta ins mostly fluid inclusions of irregular shape and highly variable liquid-to-vapor ra t ios within individual healed microfrac tures (planes of secondary inclusions) or within a group of proximal primary inclusions in a growth zone, t h e r a r e , i so la ted inclusions showing consistent liquid-to-vapor volumetr ic ra t ios a r e typically t h e smal les t inclusions. As explained l a t e r , more consis tent liquid-to-vapor ra t ios among t h e smal les t inclusions might b e a predic table consequence of t h e healing process. Homogenization t empera tu re s col lec ted f rom appropr ia te inclusions in th is t ype of qua r t z usually a r e <200°C.

Some qua r t z crys ta ls m a y conta in growth zones defined by fluid inclusions exhibit ing consistent liquid- to-vapor volumetr ic phase r a t i o s a s demonst ra ted in Figures 5.1H-5.IL and Figures 5.IP and 5.14. These t h r e e s e t s of f igures show t h r e e d i f ferent examples of growth banding in qua r t z a t low power, and each s e t has one o r two en la rgemen t s of individual a r e a s within growth zones with inclusions of consistent liquid-to- vapor ratios. F igure 5.IG shows a plane of secondary inclusions, a l l with similar liquid-to-vapor phase ratios. In growth zones o r planes where consistent liquid-to-vapor volumetr ic phase ra t ios predominate, measured homogenization t empera tu re s within a res t r ic ted a r e a of a growth zone or within a n individual plane will vary mostly within a 10-20°C range, typically around a median t empera tu re 2 2 3 0 ' ~ . The f la t , irregularly shaped primary inclusions in Figure 11 homogenize a t 2 3 0 i 1 0 ' ~ . The irregularly shaped, though more three-dimensional, inclusions in Figures 5.IK and 5.1L homogenize a t % 2 5 0 i 1 0 ~ ~ , a s do t h e secondary inclusions in Figure 5.IG (~250i5 'C) . The smooth-surfaced equan t t o negat ive crystal- shaped inclusions in Figure 5.14 homogenize a t Q270i 10°C.

The t ex tu re s discussed so far should enable a fluid inclusionist with minimal exper ience t o readily distinguish qua r t z t h a t conta ins inclusions t h a t will probably homogenize a t t empera tu re s 5 2 0 0 ' ~ f rom quar tz t h a t conta ins inclusions t h a t probably will homogenize a t 2230 '~ ; q u a r t z t h a t contains inclusions which homogenize between ~ 2 0 0 - 2 3 0 ~ C will probably be highly irregular in shape but will have more than a r a r e plane o r a r e a within a growth zone containing consistent liquid-to-vapor volumetr ic ratios. With a bit more exper ience , a fluid inclusionist will be able t o recognize inclusions which will homogenize consistently around 2 3 0 ' ~ versus inclusions which will homogenize around 270°c, based on the shapes and surfaces of t h e inclusions. Of course, during t h e

course of a fluid-inclusion study, such predictions should be suppor ted by homogenization t empera tu re s col lec ted f rom appropr ia te inclusions with a heat ing s tage , a s except ions t o t h e above generali t ies a r e ce r t a in t o occur. The general predictabil i ty of homogenization t empera tu re s based on consistency of liquid-to-vapor volumetr ic ratios, shape, and surface smoothness of inclusions may r e f l ec t t h e t empera tu re dependence of t h e kinetics of dissolution and reprecip i ta t ion of qua r t z during format ion and "maturation" of inclusions. The higher the t empera tu re , t h e f a s t e r t h e healing process, and thus t h e g rea t e r potent ia l fo r achieving the minimum su r face free-energy s t a t e (negative crys ta l morphology); t h e lower t h e tempera ture , t h e slower t h e healing process, and, thus, t h e g rea t e r potential fo r nucleation of a vapor phase prior t o complete sealing of a n individual fluid inclusion. This would also explain why smal l inclusions ( 5 3 i ~ m in longest dimension) in growth zones exhibit ing variable liquid- to-vapor ra t ios yield more consistent homogenization tempera tures . If t he se smal l inclusions a r e "pinched- off" in to s e p a r a t e en t i t i e s ear ly in t h e healing process, by v i r tue of being smal l they a r e less likely t o "neck" further.

F rom the above discussion, i t should be c lear t h a t vapor-rich inclusions can and do result from t h e healing process and in many cases may therefore not be indica t ive of boiling. Fur thermore , we emphasize t h a t boiling may have accompanied format ion of any t ex tu re s shown in Figures 5.1A-5.IL, and 5.lP-5.14, but t h a t no fluid-inclusion evidence fo r boiling was recorded in t hese examples. Figures 5.IM and 5.1N a r e photomicrographs of a qua r t z crys ta l from a vein col lec ted a t t h e nor th end of t h e Creede , Colorado mining dis t r ic t at a n elevation of 12,450 f t (ore is a t a n elevation below 11,000 ft). These inclusions demons t r a t e excel lent evidence for boiling in t he hydrothermal sys tem a t l ea s t once during growth of t h e qua r t z c rys t a l a t t h e locali ty where t h e crys ta l was growing. Figure 5.1M shows a well-defined growth zone in a qua r t z crystal , and Figure 5.1N is a n en l a rgemen t of a portion of th is growth zone showing t h a t i t is defined by (primary) fluid inclusions. Within t h e growth zone shown in Figure 5.1N a r e liquid-rich inclusions coexisting with vapor-rich inclusions. There a r e severa l impor t an t f ea tu re s t h a t indica te t h a t t h e vapor-rich inclusions result f rom en t r apmen t of a vapor phase during format ion of t h e qua r t z and a r e not a resul t of l a t e r healing (necking) processes: (1) most impor tant ly is t h e absence of one-phase, liquid-filled inclusions; (2) a l l of t h e liquid-rich inclusions show consis tent liquid-to-vapor ra t ios and yield consistent homogenization tempera tures ; and ( 3 ) al l of t h e vapor- r ich inclusions a r e la rger than t h e liquid-rich inclusions. Certainly, if t h e vapor-rich inclusions resulted f rom necking of la rger inclusions a f t e r nucleation of a vapor phase, variable liquid-to-vapor volumetr ic ra t ios among t h e liquid-rich inclusions might result ; t he se character is t ics a r e not present in Figure 5.1N. The f a c t t h a t t h e vapor-rich inclusions a r e a l l of similar s ize probably r e f l ec t s t h e s table s ize of t h e vapor bubbles (which should be of a cer ta in d iameter a t some given P, T condition) in t he boiling fluid. Also, i t might be expected t h a t some

R. J . BODNAR, T. J. REYNOLDS, & C. A. KUEHN 79

imperfections would be to ta l ly filled by liquid (forming t h e liquid-rich inclusions) and t h a t vapor bubbles might adhere t o o ther imperfections already containing some liquid. As is evident in t h e photomicrograph in Figure 5.IN, some vapor-rich inclusions do contain some liquid, most probably t rapped during format ion of the vapor-rich inclusions, which would result in incorrect ( too high) homogenization temperatures.

Vapor-rich inclusions can result from continued healing of microf r ac tu res a f t e r nucleation of vapor bubbles a s described above. If a sample contains many such microfractures crisscrossing to form a dense ar ray, a novice inclusionist could easily misinterpret t h e t ex tu re and conclude boiling. Definitive evidence t h a t a vapor phase was truly present in the hydrothermal system would be a healed microfracture containing a l l - vapor-rich inclusions a s shown in Figure 5.1 0,

- Quar tz from deeper environments (e.g., a t margins of batholiths, or around and within cupolas of batholiths, or f rom the surrounding ter ra ins undergoing greenschist t o amphibolite grade metamorphism) is character ized by the presence of abundant (millions/mm3) healed microfractures defined by small ( 5 5 ~ typically) fluid inclusions resulting in crisscrossing, sweeping, wispy, microscopic textures. These numerous secondary inclusions give "bull quartz" t h e diagnostic white color commonly observed on t h e outcrop or in hand specimens. Individual inclusions may contain only one dense, liquid phase a t room conditions o r may contain a number of phases; volati le components ( C 0 2 , CHq, N2) a t pressures g rea te r than 1 a t m a r e also typically present.

Growth zoning is not commonly observed in quar tz crys ta ls formed a t deeper levels i n the crust , but does occur. Such quar tz is distinguished from growth-banded qua r t z formed in the epi thermal environment by any one of the following possible observations: ( I ) a s t h e deep quar tz forms a t higher pressures and potentially higher temperatures , most inclusions in t h e growth zones will be smooth-surfaced and equant in shape, and most will contain e i ther a single, dense, liquid phase, or t w o fluid phases which homogenize consistently a t temperatures 5 270°C; or (2) the inclusions will contain gases under pressure easily de tec t ed by a crushing study; or ( 3 ) t he inclusions will yield very low melting points (< -20'~). Such observations contras t markedly with f ea tu res of fluid inclusions in epi thermal quar tz discussed above.

Figure 5.1R is a photomicrograph of qua r t z thought t o have originally formed a t epi thermal conditions which has been overprinted by a metamorphic even t (lower greenschist grade?) from a prospect in British Columbia, Canada. Remnant growth banding in t h e qua r t z is evident, a s a r e the ubiquitous crisscrossing microfractures defined by small C02-bear ing inclusions. Fluid inclusions defining the growth zones in Figure 5.1R display typical epithermal character is t ics , most similar t o the inclusions shown in Figures 5.IE, 5.IF, 5.IJ, and 5.IK. Whether or no t t he original epi thermal inclusions s t i l l completely preserve information regarding t h e epi thermal hydrothermal system a f t e r a metamorphic event i s debatable; more laboratory exper iments a r e necessary t o assess t h e conditions a t

which inclusions in qua r t z would begin t o be altered. In summary, petrographic examination of fluid

inclusions in qua r t z can yield valuable information concerning format ion conditions, and one need not necessarily be a n experienced fluid inclusionist t o recognize t h e important character is t ics discussed above. However, we strongly recommend t h a t a l l fluid inclusionists allow ample t ime for careful reconnaissance surveys prior t o init iating the rmomet r i c studies. Unfortunately, many fluid inclusionists tend t o e l iminate completely or spend too l i t t l e t ime on petrographic surveys. A reconnaissance petrographic examination should include many carefully se lec ted samples with consideration of the complexities t h a t may ar ise from the e f f e c t s of t ime t o der ive the maximum amount of information possible from t h e inclusions and mineral textures.

IDENTIFICATION O F FLUID INCLUSIONS TRAPPED FROM BOILING SOLUTIONS

A commonly reported f e a t u r e in studies of t h e epi thermal environment is t h e occurrence of boiling during one or more s t ages of mineral deposition. This is not surprising, a s boiling is a predictable consequence for fluids circulating a t t h e relatively high-temperature, low-pressure conditions existing in this environment. Moreover, because boiling is a potential mechanism for mineral deposition (Drummond, 1981; Drummond and Ohmoto, 1985), fluid-inclusion evidence for boiling has been proposed a s an exploration tool in t h e search for epi thermal precious-metal deposits (Kamilli and Ohmoto, 1977). In th is section, we describe various techniques for identifying fluid inclusions t h a t were trapped in a boiling system.

The coexistence of liquid-rich and vapor-rich inclusions is t h e single, most of ten c i t ed evidence in support of en t r apmen t from boiling fluids. As shown in t h e remainder of this section, this is probably also t h e mos t important l ine of evidence t h a t a fluid inclusionist may be ab le t o repor t conclusively, but severa l precautions a r e necessary. First , t h e sample must actually contain liquid-rich and vapor-rich inclusions. Because fluid inclusions a r e three- dimensional objects which a r e being observed in only two dimensions, inclusions having consistent liquid-to- vapor ratios may appear t o have variable phase ratios, a s shown in Figure 5.2. This problem is easily overcome by tes t ing the inclusions on t h e heat ing s t age because, if t he variable phase ra t ios a r e appa ren t and no t real , t h e inclusions will a l l exhibit similar modes and t empera tu res of homogenization. Second, t h e vapor-rich inclusions must result f rom en t rapmen t of a vapor phase and not from necking. As noted above, vapor-rich inclusions which result from the healing process may be distinguished from those formed by en t r apmen t of vapor by petrographic observations. Finally, t he re must be definit ive evidence t h a t t h e liquid-rich and vapor-rich inclusions a r e contemporaneous, and th is evidence is provided when both types of inclusions occur within an individual growth zone a s i l lus t ra ted in Figures 5.1M and 5.1N.

10 pm A Polished Surface H

e::l-jc::.o @.rn.@Op) 1 focus

Figure 5.2. (tog) Apparent liquid-to-vapor ratios at 25 C of numerous fluid inclusions as seen through the microscope. All fluid inclusions shown contain 25 volume-percent vapor. Inclusions numbered 1-10 along A-A' correspond to inclusions 1-10 below. (bottom) Cross-sectional view of inclusions 1- 10 shown above.

As noted in t h e previous section, addit ional def in i t ive evidence t h a t a vapor phase exis ted in a hydrothermal sys tem would be a healed mic ro f r ac tu re containing all vapor-rich inclusions (c.f., F igure 5.1 0) a s a f r ac tu re containing a l l vapor-rich inclusions could n o t resul t f rom necking alone. P re sence of a p lane of vapor-rich secondary inclusions means only t h a t boiling occurred somewhere in t h e system, n o t necessari ly a t t h e point of sample collection and perhaps a s much a s thousands of f e e t below. However, in a hydrothermal sys tem which does not i n t e r sec t a wa te r table , liquid wa te r is potentially always present, and, thus, a m e r e p lane of vapor-rich secondary inclusions may indeed indica te boiling a t t h e sample location. Thus, i t m a y be possible t o m a p zones of boiling conditions by recording re la t ive abundances of vapor-rich inclusions (regardless of the i r origin, a s long a s t hey do not resul t f rom necking) t o indica te degrees of boiling, ranging f rom minor boiling t o flashing. More r e sea rch i s required t o confirm this possibility; exper imenta l s tudies a s well a s those conducted in na tura l s e t t i ngs (ac t ive geothermal systems) a r e likely t o be fruitful .

Thus far , we have discussed t h e cha rac t e r i s t i c s of inclusions f rom t h e epi thermal environment in qual i ta t ive terms, refer r ing t o t h e inclusions simply a s liquid-rich and vapor-rich. This qual i ta t ive description may, however, be quantified, a s i t i s possible t o ca lcula te t h e phase ratios of inclusions t rapped at various PTX conditions using available phase equilibrium and PVT data. Although th is informat ion

i s generally not required fo r most fluid-inclusion studies, i t is never theless a worthwhile exerc ise because i t provides t h e fluid inclusionist with a n understanding of t h e types of inclusions he or she might expec t t o find, and helps e l imina t e obviously anomalous inclusions f rom fu r the r study. The calculations a r e particularly valuable when t ry ing t o decide if a vapor-rich inclusion might have t rapped all (or a lmost all) vapor o r a mixture of liquid and vapor.

Sample Calcula t ion I

P red ic t t h e room t empera tu re ( 2 5 ' ~ ) phase ra t ios of t h e liquid-rich and vapor-rich inclusions t rapped in a boiling, 5 wt.-% NaCl fluid a t 250 '~ . Assume t h a t e a c h inclusion t r aps only a single phase and not mixtures of liquid and vapor. The following da t a , f rom Khaibullin and Borisov (19661, P o t t e r and Brown (19771, and Keenan et al. (19781, a r e required

C n p o s i t i o n o f l i q u i d phase = 5 wt.-%NaCl

Canposi t i o n of vapor phase = 0 wt .-% NaCl

D e n s i t y of l i q u i d phase a t t r a p p i n g = 0.840 g / c m3

D e n s i t y of vapor phase a t t r a p p i n g = 0.019 g / c m3

D e n s i t y of 5 wt.-% NaCl l i q u i d a t 25% = 1.030 g / c m3

D e n s i t y of l i q u i d H20 a t 25% = 0.997 g / c m3

A t t h e t rapping conditions (250°c), t he liquid inclusion i s filled wit a 5 wt.-% NaCl liquid with a 5 density of 0.840 g/cm . A t 25OC, th is inclusion s t i l l conta ins a 5 wt.-% NaCl liquid, but now with a density of 1.030 g/cm3, a s a resul t of liquid contrac t ion during cooling. The remainder of t h e inclusion conta ins a vapor bubble t h a t is essentially a vacuum. The percent dec rease in t h e volume occupied by t h e liquid, which equals t h e volume-percent vapor, i s t he re fo re

Volune-% vapor

Thus, t h e inclusion t h a t t r apped only t h e liquid phase f rom a boiling 5 wt.-% NaCl fluid a t 2 5 0 ' ~ will 6 conta in a t 25 C a liquid occupying 81.6 volume pe rcen t of t h e inclusion a n d a vapor bubble occupying 18.4 volume percent.

Similarly, fo r t h e inclusion t h a t t r aps only t h e vapor phase

R. J. BODNAR, T. J. REYNOLDS, & C. A. KUEHN 81

V o l m - % vapor

This inclusion will conta in 98.1 volume-percent vapor and 1.9 volume-percent liquid a t 25OC.

Following t h e calculation procedure outl ined above, t h e room-tempera ture phase ra t ios of pure H 2 0 and 2, 3.5, 5, and 10 wt.-% NaCl inclusions t rapped along t h e liquid-vapor curve have been ca lcula ted and t h e results a r e l isted in Table 5.1. The room t empera tu re appearance of both t h e liquid-rich and vapor-rich inclusions t rapped f rom a boiling 5 wt.-% NaCl solution a r e a lso shown on Figure 5.3.

If inclusions a r e t rapped f rom boiling fluids, and if t h e inclusions t r a p only one or t h e o the r of t h e t w o fluids present, but not mixtures of t h e two, t h e homogenization t empera tu re s a r e t h e s a m e a s t h e trapping tempera tures . Fur thermore , t h e liquid-rich a n d vapor-rich inclusions must homogenize a t t h e s a m e tempera ture . We know, however, both f rom empir ica l observations and exper imenta l studies, t h a t t h e vapor- r ich inclusions a lmost always t r a p s o m e liquid along with t h e vapor, whereas liquid inclusions a lmos t never t r a p any vapor along with t h e liquid (Bodnar e t al., 1985; Robertson, 1968). This being t h e case, homogenization t empera tu re s of liquid-rich and vapor- r ich inclusions t rapped f rom boiling fluids might be distributed a s shown on Figure5.4, with minor variations.

Homogenization t empera tu re s of liquid-rich inclusions should a l l fal l within a narrow range, and, moreover, this range represents t h e a c t u a l range over which boiling occurred. A f e w s c a t t e r e d t empera tu re s above this value may occur and these r ep re sen t t h e f e w inclusions t h a t have leaked o r (less likely) t rapped some vapor, but a r e st i l l liquid rich. Homogenization t empera tu re s of vapor-rich inclusions* will be higher t han those of t h e coexisting liquid-rich inclusions and will be very scattered--a na tura l consequence of t h e f a c t t h a t t h e vapor-rich inclusions nearly a lways t r a p some Iiquid a n d t h a t t h e proportions of liquid a n d vapor t rapped in these inclusions will thus be variable. The maximum possible homogenization t empera tu re of t h e liquid-rich inclusions i s t h e c r i t i ca l t empera tu re corresponding t o t h e bulk composit ion of t h e mixture in t h e inclusion. Homogenization t empera tu re s of vapor-rich inclusions, on t h e o the r hand, may be g rea t e r than t h e c r i t i ca l t empera tu re fo r t h e bulk composition in t h e inclusion, and may extend t o t h e maximum tempera tu re on t h e solvus f o r t h a t composition (Bodnar e t al., 1985). If a vapor-rich inclusion t r aps so much liquid t h a t t h e bulk inclusion density is g rea t e r than t h e cr i t ica l density, t h e inclusion will behave like a liquid-rich inclusion. T h a t is, during heating, t h e vapor bubble will shrink and disappear a t some t empera tu re between t h e boiling t empera tu re and t h e cr i t ica l tempera ture . Many of t h e anomalously high homogenization t empera tu re s of

Table 5.1--Calculated volume-percent vapor at 25'~ of H20-NaC1 fluid inclusions trapped from boiling solutions

Salinity Temperature (OC) (wt. % ) 150 200 250 300

L = fluid inclusions that trapped all liquid. V = fluid inclusions that trapped all vapor.

Figure 5.3. Room temperature (25Oc) pnase relations of H20-NaC1 fluid inclusions having a salinity of 5 weight-percent NaCl and homogenization temperatures from 1 5 0 ~ ~ to 3 0 0 ~ ~ . Also shown are the room temperature phase relations of inclusions that trapped the vapor phase that would have been in equilibrium with the liquid phase a t each temperature. Black corresponds to the vapor phase and the unshaded portion represents the liquid phase.

liquid-rich inclusions in epi thermal deposits perhaps a r e a resul t of th is somewha t special ca se of heterogeneous ent rapment .

*Homogenization t empera tu re s of vapor-rich inclusions depic ted in Figure 5.4 represent t h e t empera tu re s t h a t would be observed if such d a t a were obtainable. In prac t ice , homogenization t empera tu re s of vapor-rich inclusions can rare ly be measured because t h e inclusion geomet ry usually prevents t h e fluid inclusionist f rom viewing t h e final disappearance

0.042 g + 0.018 g d e n s i t y =

1 c m 3

Homogenization Temperature -+

Figure 5.4. Schematic representation of the manner in which liquid-rich and vapor-rich inclusions trapped from a boiling solution would be distributed assuming the inclusions trapped either all liquid or mixtures of liquid and vapor, but never only vapor.

of liquid, t h a t is, assuming a dist inct liquid phase is observable a t all, and the precision of t hese measu remen t s is usually very poor (*10-25OC). This problem becomes less severe a s (1) t h e inclusion geometry becomes m o r e irregular, (2) t h e volume- percent liquid in t h e inclusion increases, and (3) t h e homogenization t empera tu re nears t h e cr i t ica l tempera ture .

This qual i ta t ive description of t h e manner in which homogenization t empera tu re s of inclusions t rapped in a boiling sys t em should be distributed may be quantified? provided t h a t t h e necessary PVT and phase-equilibria d a t a f o r t h e fluid of in teres t a r e available. Moreover, t h e salinit ies of inclusions t rapping mixtures of liquid and vapor may be ca lcula ted , and t h e variation in salinity with homogenization t empera tu re predicted.

Sample Calcula t ion 2

Using d a t a provided in Sample Calculation 1 above, ca lcula te t h e bulk-inclusion density and composition of a n inclusion t h a t t r aps 95 volume- percent vapor and 5 volume-percent liquid f rom a boiling, 5 wt.-% NaCl sol t ion a t 250 '~ . Assuming a n inclusion volume of 1 cmf: t h e masses of liquid and of vapor t r apped in t h e inclusion a r e

mass of l i q u i d = 0.05 c m3 x 0.840 g / c m3

mass of vapor = 0.95 c m 3 x 0.019 g / c m 3

Of t h e 0.042 g of liquid t r apped in t h e inclusion, 5 weight pe rcen t is NaCl and 95 weight percent is H20. Assume t h a t t h e t rapped vapor i s pure H20 . The mass of NaCl t rapped in t h e inclusion is

~ m s s NaCl = 0.05 x 0.042 g

Therefore, t h e composit ion of t h e inclusion i s

0.0021 g s a l i n i t y = x 100

0.042 g + 0.018 g

Thus, th is inclusion, which t rapped 5 volume-percent liquid having a composit ion of 5 wt.-% NaCl and 95 volume-percent vapor having a composition of pure H 2 0 , has a bulk density of 0.060 g/cm3 and a salinity of 3.50 wt.-% NaCl. Although t h e e x a c t homogeniza- t ion t empera tu re of th is inclusion cannot be de t e rmined owing t o lack of PVTX d a t a in th is region, we can de t e rmine t h a t t h e homogenization t empera tu re will be g rea t e r than t h e cr i t ica l t empera tu re of a 3.5 wt.-% NaCl fluid, i.e., g r ea t e r t han 410°C (Khaibullin and Borisov, 1966).

Following t h e procedure outl ined in Sample Calcula t ion 2 and using t h e d a t a of Khaibullin and Borisov (19661, t h e composit ions and homogenization t empera tu re s of inclusions t rapping various proportions of liquid and vapor f rom a boiling, 5 wt.-% NaCl solution a t 250°C have been ca lcula ted and a r e shown o n Figure 5.5. Homogenization t empera tu re s of liquid- r ich inclusions may vary f rom t h e ac tua l t rapping t empera tu re (250 '~ ) t o t h e cr i t ica l t empera tu re fo r t h e bulk-fluid composition ( 4 1 7 ~ ~ 1 , but may not exceed th is value. Homogenization t empera tu re s of vapor-rich inclusions may extend t o much higher values, but t h e e x a c t t empera tu re s cannot be predic ted because of lack of PVTX d a t a in th is region. The salinity may be any value between t h e liquid-phase composit ion (5 wt.-% NaCI) and t h e vapor-phase composit ion (pure H20).

If a group of apparent ly contemporaneous fluid inclusions exhibits variable liquid-to-vapor ratios, i t is common p rac t i ce t o not col lec t homogenization t e m p e r a t u r e and salinity d a t a f rom these inclusions a s t hey have obviously t rapped mixtures of liquid and vapor (or have leaked or necked). While i t is t r u e t h a t t h e measured values do not represent t h e ac tua l t rapping t empera tu re s or composit ions of t h e fluids in

Volume Percent Vapor Trapped

Figure 5.5 Calculated variation in inclusion composition and homogenization temperature as a function of the volume-percent vapor trapped from a hiling, 5 wt.4 NaCl solution at 250%.

equilibrium a t trapping, they may provide addit ional evidence t h a t t h e inclusions were t rapped in a boiling sys tem, owing t o t h e sys t ema t i c and predic table manner in which homogenization t empera tu re var ies a s a function of salinity, a s shown on Figure 5.6.

In summary, f rom a prac t icable standpoint, t h e occurrence of coexisting, contemporaneously formed liquid-rich and vapor-rich inclusions in a given sample is t h e bes t evidence t h a t could normally be observed t o indica te t h a t t h e inclusions were trapped during boiling. However, if homogenization t empera tu re and salinity d a t a can be obtained f rom such inclusions, addit ional evidence supporting a boiling in tepre ta t ion may be determined, owing t o t h e sys t ema t i c and predic table manner in which these values must vary if t h e inclusions t r apped mixtures of liquid and vapor. I t is c l ea r t h a t our understanding of t h e trapping mechanisms and behavior of inclusions t rapped f rom boiling fluids must be improved, and th is will require addit ional exper imenta l s tudies a s well a s de ta i led s tudies of inclusions f rom na tu ra l sys tems where boiling is known t o have occurred.

Salinity (wt % NaCI)

Figure 5.6. Relationship between homogenization temperature and salinity of inclusions trapping various proportions of liquid and vapor from a boiling, 5 wt.-% NaCl solution at 2 5 0 ~ ~ .

IDENTIFICATION O F GASES IN FLUID INCLUSIONS FROM THE EPITHERMAL ENVIRONMENT

The presence of gases in fluid inclusions f rom t h e epi thermal environment has been repor ted by numerous workers, a n d these volati le components have been shown t o play a n impor tant ro le in controll ing solution chemistry. Moreover, t h e presence of gases can significantly a f f e c t t h e pressure, and, therefore , t h e depth, a t which boiling may commence in t hese systems. In th is sec t ion we describe t h e techniques f o r recognizing gases in fluid inclusions and explain why t h e presence of gases may not have been recognized in many fluid-inclusion s tudies of epi thermal deposits.

The predominant volati le component in most fluid inclusions f rom t h e epi thermal environment i s COZ, and we will base our discussion on t h e PVTX proper t ies of t h e H 2 0 - C 0 2 system. Phase re la t ions in t h e sys t em H 0 C O a r e well known over a wide r ange of PTX condftions K l ~ i s and Golding, 1963; Todheide and Franck, 1963; Takenouchi and Kennedy, 1964; Malinin and Kurovskaya, 1975). An impor tant f e a t u r e

of this sys tem with r e spec t t o epi thermal deposits is t h e two-phase field in which H20-rich fluids coexist with more C02-r ich fluids (Fig. 5.7). This f ield of immiscibil i ty encompasses only a relatively small portion of PTX space a t t empera tu re s g rea t e r than about 300°c, but expands considerably a s t empera tu re decreases. Over t h e t e m p e r a t u r e range in which most epi thermal deposi t s formed ( I O O ~ C - ~ ~ ~ ~ C ) , only very H20-rich o r very C02-r ich composit ions do not l ie in t h e two-phase field, even up t o pressures of severa l kilobars, excep t fo r t h e sma l l one-phase "vapor" field t h a t spans t h e e n t i r e composit ion range a t very low pressures (Fig. 5.7).

The presence of C 0 2 in fluid inclusions may be recognized by ( I ) t h e occu r rence of three-phase (liquid H 2 0 , liquid C O , and vapor C 0 2 ) inclusions a t room tempera ture , (2f) expansion of t h e vapor bubble when

Figure 5.7. Isotherms showing the mmposit ions of coexisting liquid and vapor phases in the system H20-C02 and H20M2;NaC1. Also shown are the 2 5 O ~ phase relations of H20- C02 fluid inclusions with a bulk composl- tion of 30 mole-% C02 trapped at various temperatures and pressures represented by points A-F. The innermost phase in each inclusion (black) represents a C02-rich vapor, the outermost phase (white) represents an H20-rich liquid, and the intermediate phase (stippled pattern) represents a C02;rich liquid. W m temperature phase relations of these inclusions were calculated as described in the text.

t h e inclusions a r e opened by crushing t h e sample in oil, (3) t h e nucleation of t h e C 0 2 . c l a t h r a t e compound ( C 0 2 , ' 5 314 H 2 0 ) during coollng of a two-phase inclusion f rom room t empera tu re t o lower tempera tures , and (4) microanalysis of fluids released f rom inclusions. The t empera tu re and pressure of t rapping and t h e amoun t of C 0 2 required t o genera te three-phase inclusions a t room tempera ture , or t o produce inclusions which will nuclea te a liquid C 0 2 phase during cooling, or show expansion of t h e vapor bubble during crushing t e s t s may be ca lcula ted f rom avai lable phase equilibrium and PVT d a t a fo r t h e H 2 0 - C 0 2 system. Because fluid inclusions a r e constant-volume systems, t h e only information required t o ca lcula te t h e room t empera tu re phase r a t i o s and t empera tu re s and pressures of various phase changes in t h e inclusions i s t h e bulk density of t h e inclusion. The bulk density of a n H20-C02 fluid a t a given t empera tu re and pressure of t rapping may be determined f rom a modified Rediich-Kwong equation of s t a t e (Connolly and Bodnar, 1983).

Once t h e bulk-inclusion density is known, t h e re la t ive amounts of t h e H20-rich hase and t h e C 0 2 - g r ich phase(s) in t h e inclusion a t 25 C may be obtained by determining a t which pressure ( a t T = 2 5 ' ~ ) t he necessary mass-balance and phase-equilibria requirements imposed by t h e inclusion density and composit ion a r e satisfied. This pressure, obtained using i t e r a t i ve computa t ion techniques shown on t h e flow c h a r t in Figure 5.8, is unique fo r a given inclusion composit ion and density. The init ial guess a t t h e in ternal pressure in t h e inclusion a t 25OC is 64 bars, corresponding t o a n inclusion containing both liquid and vapor C 0 2 a t 2 5 ' ~ . T h a t is, 64 bars corresponds t o t h e equilibrium vapor pressure of C 0 2 a t 25 '~ . A f t e r ca lcula t ing t h e amoun t of H 2 0 in t h e inclusion f rom t h e e s t ima ted bulk composition, t h e amount of C 0 2 t h a t will dissolve in t h a t mass of H 2 0 a t 25OC and 64 bars is de termined using t h e d a t a of Dodds e t al. (1956). Then, t h e volume occupied by this H20-rich phase is de termined using density d a t a fo r H 2 0 - C 0 2 solutions a t low t empera tu re s f rom Parkinson and d e Nevers (1969).

The amount of "free" C 0 2 in t h e inclusion i s t h e t o t a l amoun t of C 0 2 , again obta ined f rom t h e known bulk composition, less t h a t amoun t of C O dissolved in t h e H20-rich phase." This amoun t of 2 0 2 mus t fill t h a t portion of t h e inclusion no t occupied by the H 0 r ich phase. The volume and mass of t h e "free" 20; phase may be obta ined f rom mass-balance calculations, because of t h e assumption t h a t fluid inclusions represent constant-volume, constant-mass, chemical systems; t hese values provide t h e density of t h e "free" C O phase in t h e inclusion. If this ca lcula ted bulk $0 density i s g r e a t e r t han t h e density of C 0 2 vapor on tfie C 0 2 liquid-vapor curve a t 25'C and less than t h e dens i ty of liquid C 0 2 , a t 25OC and vapor sa tura t ion , t h e inclusion will conta in both liquid and vapor C O a t 2 5 ' ~ and have a n in ternal pressure of 64 bars. l%e r e l a t i ve amounts of liquid and vapor C 0 2 in such a n inclusion a r e obtained f rom mass balance calculations.

+In t h e calculations of Burruss (1981a,b), Hollister and

Compute Inclusion Bulk Density as F (P, T, Xco,) Using Redlich-

Kwong Equation of State

I Compute Relative Volume 1 Proportions of Aqueous Phase and

"Free" CO, Phase

* I Compute Density of "Free" Co, I I

co2 (V) < P < CO, (L)

Density of "Free" COP

v ) Compute Density of "Free" C02 1 I at 25°C and P(New) I I Compute P(lnt) as Function of I I Density of "Free" CO, I

v Compute Phase

Yes + Ratios, Th (CO,), Tm ( H Y ~ )

A

Density, Phase Ratios,

Figure 5.5. Flow cha r t of the computer program used t o ca l cu l a t e room temperature phase r a t i o s and other proper t ies of H20-C02 f l u i d inclusions.

Burruss (19761, and Ramboz e t al. (19821, the amount of C 0 2 dissolved in the H 0 phase is neglected, i.e., considered to be zero. ?'his assumption does not greatly a f fec t the calculated phase relations and temperatures of phase changes for most inclusions in the H20-C02 system. However, for some inclusions, particularly those with low C 0 2 contents or with internal pressures greater than a few tens of bars, neglecting the amount of C 0 2 dissolved in H 2 0 greatly affects the calculated low-temperature properties (Swanenberg, 1980). The amount of H 2 0 that will dissolve in C 0 2 a t 2 5 ' ~ and pressures greater than the equilibrium vapor pressure of pure H 2 0 a t this temperature ( 0.03 bars; Keenan e t al., 1978) is approximately an order of magnitude less than the amount of CO that will dissolve in H 2 0 a t these same conditions. ~ z e r e f o r e , the C 0 2 phases were assumed to be pure in the calculation of room temperature phase ratios of inclusions. This assumption introduces negligible errors into the calculations.

Sample Calculation 3

Assume tha t H20-CO fluid inclusions a re trapped in the liquid-vapor field a t 2 5 0 ' ~ and 700

the density of the "free" C02 . in the H20-rich inclusion? What a r e the phase r a t ~ o s of this inclusion a t 25Oc?

The total mass of fluid (H20 + C02) in the inclusion is

Tota l m s s = 100 c m3 x 0.831 g / c m3 = 83.1 g

The mass of H 2 0 in the inclusion is

(0.916 moles x 18 g/mole) x 83.1 g

(0.916 moles + (0.084 moles x 18 g / m l e ) x 44 g/mole)

The number of moles of H 2 0 in the inclusion is

67.88 g Moles H p = = 3.77 moles

18 g/mole

The total mass of C 0 2 in the inclusion is

The total number of moles of C 0 2 in the inclusion is

15.22 g Moles 032 = = 0.346 moles

44 g / m l e

Now, for the initial i teration in the calculation procedure, assume that a t 2 5 ' ~ the inclusion contains three phases--liquid H 2 0 and liquid and vapor C02. At 2 5 ' ~ , the equilibrium vapor pressure of C 0 2 IS, 6 4 bars (Quinn and Jones, 1936). Thus, our i n ~ t l a l assumption is that the internal pressure of the inclusion is 64 bars. At 2 5 ' ~ and 64 bars, liquid H 0 has a density of 0.9998 g/cm3 (Keenan e t al., 197b. Thus, the volume of liquid H 2 0 in the inclusion is

Now, the aqueous phase is not pure H 0 but, rather, is a C02-saturated, H20-rich phase. ?he solubility of C 0 2 in H 2 0 a t 2 5 ' ~ and 64 bars is 2.433 mole percent (Fig. 5.9). So, the total number of moles of C 0 2 dissolved in the H 2 0 phase in our model inclusion is

0.02433 x 3.77 moles Moles a2 =

1.0 - 0.02433

= 0.094 moles CD2

C02 Solubility in H20 (Mole Percent) as a Function

of Pressure at 25°C

1 0 1 .O 2.0 3.0

CO, Solubility (Mole Percent)

Figure 5.9. S o l u b i l i t y of C02 i n H20 as a function of pressure a t 25%.

R. J. BODNAR, T. J. REYNOLDS, & C. A. KUEHN 87

The volume of C 0 2 in the C02-satura ted aqueous solution a t 2 5 ' ~ and 64 bars assumlng a molal volume 1 of C 0 2 in solution of 29 cm /mole (Wiebe and Gaddy, 19391, 1s

Volume of C 0 2 in solution

67.88 g = (0.094 moles x 29 c m3/mole) x

1000 g

The to t a l volume of the C02-satura ted aqueous liquid phase is

Volune aqueous phase = 67.89 c m3 + 0.185 c m3

The mass of C 0 2 remaining ("free" C 0 2 ) in the inclusion is

= (0.346 moles - 0.094 moles ) x 44 g / m l e

The volume occupied by "free" C 0 2 is

"01- " f r e e " 032 = 100 c m3 - 68.075 c m3

The bulk density of "free" C 0 2 is

11.08 g D e n s i t y of " f r e e " 032 =

31.92 c m3

Because the density of ''free" C 0 2 is greater than t h e d e i t y of C 0 2 vapor a t 2 5 ' ~ and 64 bars Y (0.2375 g/cm ), and less than t h e d nsity of C O liquid 5 . a t 2 5 ' ~ and 64 bars (0.71 g/cm ) (Flg. 5.18), t h e inclusion will contain both liquid and vapor C 0 2 a t 2 5 ' ~ .

The volume of C 0 2 liquid in this inclusion is

Volume of C 0 2 liquid

11.08 g - (0.2375 g / c m3 x 31.92 c m3) -

0.479 g / c m3

= 7 .3 c m3

The volume of C 0 2 vapor is

Volune of 032 vapor = 31.9 c m3 - 7.3 c m3

= 24.6 c m3

Density (g.cm-9

F i g u r e 5.10. D e n s i t y o f l i q u i d and vapor C02 a t p h a s e e q u i l i b r i u m from OOC to the cri t lcal t e m p e r a t u r e ( 3 1 . 1 ~ ~ ) .

Therefore, assuming a to t a l inclusion volume of 100 cm3, the inclusion will contain

68.1 vol lme-percent aqueous phase 24.6 vo lune -pe rcen t (332 vapor

7 .3 v o l w - p e r c e n t 032 l i q u i d

100.0 vo lune p e r c e n t

Homogenization t empera tu res of t he C 0 2 phases ("free" C 0 2 ) in H 2 0 - C 0 2 inclusions may be es t imated using the ca lcula ted density of "free" C 0 2 and t h e known densit ies of liquid and vapor on the C 0 2 solvus (Fig. 5.10). The homogenization t empera tu re of the "free" C 0 2 phases is t h a t t empera tu re on t h e liquid- vapor curve a t which CO2, e i ther liquid or vapor, t h a t has t h e same density a s t h a t of "free" C 0 2 a t 2 5 ' ~ and a pressure equal t o the ca lcula ted internal pressure of t h e inclusion a t this temperature . Obviously, if t he inclusion has a n in ternal pressure of 64 bars and contains both liquid and vapor C 0 2 a t 25OC, the homogenization t empera tu re is between 2 5 ' ~ and t h e cr i t ica l t empera tu re (31.0'~). Similarly, if t he inclusion contains only one C 0 2 phase a t 2 5 O ~ , e i ther liquid or vapor, t he C 0 2 phases a r e already homogenized and t h e homogen iza t~on t empera tu re is less than 25OC. From Figure 5.10, de t e rmine t h e homogenization temperature , and t h e phase t o which t h e "free" C 0 2 phases will homogenize. (Answer: T h = 30.3'~, t o vapor phase).

Following the calculation procedure outlined above, t h e room temperature phase ra t ios of H 2 0 - C 0 2 inclusions t rapped a t t h e PTX conditions shown on Figure 5.1 1 have been determined. The appearance of these inclusions a t 2 5 ' ~ is shown on Figure 5.12 and the d a t a a r e l isted in Table 5.2. Also l isted in Table 5.2 a r e the ca lcula ted homogenization t empera tu res of t he C 0 2 phases, t h e dissociation

t empera tu re of t he c l a th ra t e , and t h e internal pressure a t 25 '~ . All inclusions were t rapped f rom a fluid having a bulk composition of 4 mole-percent C 0 2 . A t 450 bars (point I , Fig. 5.1 11, a l l t empera tu res f rom 2 0 0 ' ~ t o 3 0 0 ' ~ a r e in t h e one-phase field, and fluid inclusions t h a t t r a p this fluid would contain t w o phases, liquid water and C 0 2 vapor a t 2 5 ' ~ (Fig. 5.12, Table 5.2). A t 350 bars (point 2, Fig. 5.1 1) t h e fluid is also in t h e one-phase field a t 2 5 0 ' ~ and 3 0 0 ' ~ and the resulting inclusions would b e similar t o those trapped a t 450 bars. However, a t 2 0 0 ' ~ and 350 bars, a 4 mole-% composition is in t h e two-phase field, and would split in to t w o immiscible fluids. A t 2 5 ' ~ fluid inclusions of t h e H20-rich phase (96.4 mole-% H20; Table 5.2) would conta in t w o phases, liquid water and C O vapor; those of t h e C02-r ich phase (16.2 mole-% ~ ~ 8 ) would contain t h r e e phases, liquid water , and liquid and gaseous C 0 2 (Fig. 5.12). With slight heating, t h e C 0 2 phase in t h e C02-r ich inclusion would homogenize t o t h e vapor phase a t 30.9OC (Table 5.2).

I / - 1 One-Phase Field

1 . l i

I I I I I I 200"

2 4 6 8 Hz0 Mole % CO,

F i g u r e 5.11. The 200°c, 250°c, a n d 3 0 0 ~ ~ isotherms i n the l o w p r e s s u r e ( 6 0 0 b a r s ) , w a t e r - r i c h p o r t i o n o f t h e H20-C02 phase d i a g r a m . The i s o t h e r m s r e p r e s e n t t h e b o u n d a r y b e t w e e n t h e o n e - p h a s e f i e l d (up- p e r - l e f t p o r t i o n o f the f i g u r e ) a n d the two-phase f i e l d ( l o w e r - r i g h t p o r t i o n o f the f i g u r e ) a t tha t p a r t i c u l a r t e m p e r a t u r e . P o i n t s 1-5 c o r r e s p o n d t o p o i n t s 1-5 o n F i g u r e 5.12.

A t 250 bars (point 3, Fig. 5.111, a 4 mole-% C 0 2 bulk composition would be in t h e two-phase field a t 200°C and 250°C, but would be in the one-phase field a t 3 0 0 ' ~ . Inclusion pairs t rapped at 200°C and 250°C would be of t he H20-rich and C02-r ich var ie t ies when observed a t 25Oc, and t h e C O r ~ c h ~nc lus ion trapped a t ZOOOC would contain a s m a h a m o u n t of liquid C O ~ (Fig. 5.12; Table 5.2).

A t pressures of 200 and 100 bars (points 4 and 5, Fig. 5.1 I), al l t empera tu res between 200°C and 3 0 0 ' ~ a r e in t h e two-phase field for a bulk composition of 4 mole-% C 0 2 . Inclusions t h a t would result from trapping e i ther t h e H 0 rich or t h e C02-r ich phase a t these conditions woufd be liquid rich and vapor rich, respectively, and a l l would conta in only two phases, liquid water and C 0 2 vapor a t 2 5 ' ~ (Fig. 5.12, Table 5.2).

Several important f ea tu res a r e revealed by the phase ra t ios shown on Figure 5.12 and t h e d a t a listed in Table 5.2. Most importantly, three-phase inclusions a r e rarely produced a t t hese PTX trapping conditions. This in spi te of t he f a c t t h a t we chose a C 0 2 concentration (4 mole %) t h a t is much higher than normally encountered in the epi thermal environment. Had we used a more rea l is t ic C 0 2 content fo r t h e epi thermal environment, such a s 1.0 mole-% C 0 2 or less, three-phase inclusions would be completely absent. Moreover, only the vapor-rich inclusions contain th ree phases and, owing to optical problems, t he presence of liquid and vapor C 0 2 in these inclusions might easily be overlooked. Secondly, t he liquid-to-vapor r a t io of t he two-phase inclusions is very similar t o t h a t of simple H20-salt inclusions t rapped in the s a m e t empera tu re range (compare with Fig. 5.3). As a result , even though these inclusions contain significant amounts of C 0 2 , t he presence of C 0 2 would no t b e revealed by simple petrography.

The discussions above show t h a t t h e presence of C 0 2 in fluid inclusions f rom epi thermal deposits will generally not be discernible from room temperature phase relations. However, one relatively simple t e s t for COJ, or more appropriately for t he presence of noncon ensed gases, i s t h e crushing test . If a fluid inclusion contains water or a simple sa l t solution i t would have an equilibrium vapor pressure of about 0.03 bars a t 2 5 ' ~ (Keenan et al., 1978). Therefore, if we immerse t h e sample containing this inclusion in oil and break the inclusion open, t h e oil will be forced into the inclusion because t h e external pressure (%I bar) is considerably higher than t h e internal pressure of t he inclusion (%0.03 bars). If we observe this procedure under the microscope, t he bubble will appear t o collapse and disappear a t t h e ins tant t he inclusion is opened. If t h e inclusion contains some dissolved, noncondensed gases, t h e internal pressure will be greater than 0.03 bars. Under these conditions the vapor bubble will e i the r only partially collapse if 0.03 <Pint <l.O, or will expand if Pint> 1.0 bar.


By what amount will t he C 0 2 phases in t h e inclusion in Sample Calculation 3 above increase in

Table 5.2--Calculated room temperature properties of H20-C02 fluid inclusions

Trapping temperature (OC) Trapping pressure (bars) Composition (mole-fraction H20) Volume-percent aqueous liquid phase at 25'~ Volume-percent liquid C02 at 25'~ Volume-percent C02 vapor at 25'~ Homogenization temperature of C02 phases (OC) Clathrate dissociation temperature (OC) Internal pressure at 25'~ (bars)

Figure 5.12. Room temperature (25O~) phase relations of H20-C02 fluid inclusions trapped at temperatures of 200°c, 250°c, and 3 0 0 ~ ~ and pressures of 100, 200, 250, 350, and 450 bars, corresponding to points 1-5 on Figure 5.11. The innermost phase in each inclusion (black) represents a C02- rich vapor and the outer phase (white) represents an H20-rich liquid. The phase between the COZ-rich vapor and the H20-rich liquid in the inclusions trapped at 2 0 0 ~ ~ and 250 and 350 bars (stippled pattern) represents a C02-rfch liquid. Room temperature phase relations of these inclusions were calculated using the method described in the text and are listed along with calculated temperatures of various phase changes in Table 5.2.

volume (expand) when t h e inclusion i s opened by crushing t h e sample in oil? The bulk density of C 0 2 in t h e unopened inclusion i s 0.347 g/cm3, and he dens i ty 5 of C 0 2 a t 2 5 ' ~ and 1 bar is 0.00172 g/cm . Neglect t h e C 0 2 dissolved in t h e H 2 0 liquid phase which, under normal conditions, would exsolve f rom t h e H 2 0 phase when t h e inclusion is opened and contr ibute t o t h e expansion of t h e vapor bubble.

0.347 g / c m3 vo lune i n c r e a s e = = 202 t imes

0.00172 g / c m3

Thus, t he volume of t h e C 0 2 phase a f t e r crushing will be 202 t imes larger t han t h e volume occupied by t h e C 0 2 phases in t h e unopened inclusion. If t h e C 0 2 phases in t h e unopened inclusion had a d iameter of 3 microns, assuming a spherical geometry , t h e C O bubble fo rmed a f t e r crushing would have a d i ame te r o$ approximate ly 17.5 microns.

Figure 5.13A shows a qua r t z chip f rom t h e Car l in sediment-hosted disseminated-gold deposi t in t h e crushing s t age before t h e crushing t e s t was begun. During t h e crushing tes t , numerous smal l bubbles appeared a s t h e qua r t z f laked off t h e edge of t h e chip (Fig. 5.13B, C , D). These bubbles represent t h e con ten t s of fluid inclusions t h a t have expanded upon opening of t h e inclusions, indicating t h a t t h e inclusions conta ined a gas a t a n in ternal pressure g rea t e r t han ambien t a tmospher ic pressure ( 1 bar). Because of t h e smal l s i ze of t h e inclusions in t h e s ample shown in Figure 5.13, a s is typical of epi thermal deposits, individual inclusions could not be observed a s they opened; e a c h bubble in Figure 5.13 conta ins t h e con ten t s of numerous fluid inclusions. This is suggested by t h e observation t h a t during t h e crushing process s eve ra l t ens of smal l bubbles rapidly coalesced in to one large, s table bubble (Figs. 5.13C, D). O the r smal l bubbles t h a t were no t incorpora ted in to t hese larger bubbles quickly dissolved i n t o t h e oil.

The inclusions conta ined in t h e Car l in sample shown in Figure 5.13 a r e n o t typica l of inclusions f rom most epi thermal deposits because they conta in on t h e o rde r of 5-20 mole-% COZ, whereas mos t inclusions f rom bonanza-type vein deposits conta in on t h e order of a f ew t en ths of a mole-% C 0 2 (Fig. 5.14). Moreover, when crushing s tudies a r e conducted on inclusions f rom bonanza-type deposits , evidence of gases (i.e., expanding vapor bubbles) is rare ly found. The question, then, i s how much C 0 2 i s required before we will s e e t h e bubble expand during crushing? This value may b e easily ca lcula ted f rom PVTX d a t a fo r t h e H 2 0 - C 0 2 system.


Assume we have an H O - C 0 2 inclusion containing a vapor bubble (all 80 ) occupying 25 volume percent of t h e inclusion a t 2 3 0 ~ and t h a t t h e in ternal pressure is I bar. How much C 0 2 does t h e inclusion contain?

A t 2 5 ' ~ and I bar, t h e solubility of C O in H 0 is 0.15 wt.% o r 0.06 mole % (Fig. 5.9; ~ o d f s t aT., 1956). Assuming a n inclusion volume of 100 cm', t he C02- sa tu ra t ed liquid, with a densi ty of approximate ly

Figure 5.13. Crushing test on a quartz chip containing high-density, CD2-rich inclusions from the Carlin deposit.

A). Quartz chip in oil in crushing stage prior to beginnrng of crushing test. Fluid inclusions in this sample are small (cl0 micrometers) and appear as dark specks and diffuse bands where t??e iaclusions occur along fractures.

5). Initial stages of crushing test. As quartz fragments begir. to flake off the edges of the chip, numerous inclusions are opened, releasing gases into the oil (lower and upper right edges of the chip).

C). With continued crushing mre inclusions are opened and smaller bhbles representing gases released from one or a few inclusions coalesce to from larger bubbles.

Dj. Final stages of crushing test. Several large bubbles occur away from chip. The interface between crushed and uncrushed sample is compsed of a mixture of gas-saturated oil and finely ground quartz arid appears as a dark "meniscus" surrounding the uncrushed quartz.

T h e to t a l mass of C 0 2 in t h e inclusion (GO2 vapor plus C 0 2 in solution) equals

Ngawha

@ Kaweru

@ Waiotapu

Ohaki / T o t a l CD2 i n i n c l u s i o n = 0.1122 g + 0.043 g I

T h e to t a l mass of fluid in t h e inclusion ( H 2 0 plus GO2) eqirals

T o t a l f l u i d i n i n c l u s i o n = 0.1552 g + 74.6628 g

= 74.818 g

Therefore, t he bulk composition of t h e inclusion equals

0.1552 g Buik c a r p o s i t i o n =

74.818 g

= 0.207 wt.-%CO2

Acupan Bagio I = 0.047 m i a l

I Topia

1 National District I

Mole Percent GQ,

Figure 5.14. Summary of reported CC concentra- t ions of hydrothermal fluids &om several terrestr ial geothermal systems and epithermal precims-metal deposits. Compiled from data given by Hederquist an6 Henley (1985), Sommer et aL. (1985) and VYkre (1985).

0.997 g/cm3 (Parkinson and d e Nevers, 1969), has a to t a l mass of

Nhss s o l u t i o n = 10.997 g / c m3) (75 c m3)

Of this t o t a l mass, @.I5 wt.-% is C 0 2 or

!%ss 9 i n s o l u t i o n = ( .0015) (74.775 g)

The vapor bubbl (assumed t o be a l i COZ) has a density 5 . of 0.00172 g/cm ( Q u ~ n n and Jones, 1936) and contains a to t a l mass of C 0 2 of

@ass a2 i n vapor = (0.00!72 g / c m3) (25 c 2)

Thus, t h e vapor bubble of a n Ef20-C02 fluid inclusion with a bulk composition 0.047 rnolal 1 0.086 mole percent) will no t expand when t h e inclusion is crushed and observed under t h e microscope.

The minimum amount of C 0 2 required before t h e vapor bubble will expand during crushing studies ( d .1 moie percent) is within t h e range of C 0 2 concentra t ions commonijr repor ted f rom modern t e r r e s t r i a l geothermal sys tems and their fossil analogs, t h e epi thermai precious-metal deposits (Fig. 5.14). Therefore, if inclusions frorn bonanza-type epi thermal deposits a r e crushed in oil, t h e explosive re lease of gas and format ion of large vapor bubbies in t h e oil (as shown in Fig: 5.13) will no t b e observed. This is exact ly what 1s found when these t e s t s a r e made, and t h e presence of C O is never indicated by crushing t e s t s on inclusions I rom most epi ther inai precious- m e t a l deposits.

Using slightly higher GO2 contents of 0.2 and 0.5 mole percent, which is st i l l i n t h e range of repor ted values from the epi thermal environment (Fig. 5-14], t h e 25% internal pressures of inclusions t rapped a t epi thermal P-T conditions have been calcula ted and a r e shown in Figlire 5.15. For a C 0 2 content of 0.2 moie percent, t he in ternal pressures of inclusions f rom t h e epi thermal environment will be 4-6 bars a t 25'C.

Consider a n H20-CO fluid inclusion with a bulk C 0 2 concentration of 0.3 mole pe rcen t t rapped a t 300 C and 300 bars. Using t h e calculation procedure described above, this inciusion will contain t w o phases at 25'C--an aqueous liquid phase and a C 0 2 v?por bubble occupying 25.2 volume pe rcen t of t h e inclus~on. Fur thermore , t he internal pressure of this inclusion a t 25% will be 5 bars and t e GO2 vapor phase will have 9 a density of 0.010 g/cm . What will happen t o the vapor bubble in this inclusion if we crush t h e sample and open t h e inclusion in oil? From t h e calcuia ted

R. J. BODNAR, T. J. REYNOLDS, & C. A. KUEHN

density of t h e C O vapor phase in t h e unopened 23 inclusion (0.010 g/cm ) and the know? density of C 0 2

gas a t 25OC and I bar (0.00172 g/cm ), we can predict t h a t t h e vapor bubble volume a f t e r crushing will be 5.8 t imes i t s volume before crushing. If t h e vapor bubble d iameter was initially 2 microns, i t s d iameter a f t e r crushing would be 3.6 microns. As a result, explosive re lease of C O into t h e surrounding oil t o produce large, easily o8servable vapor bubbles, a s shown in Figure 5.13, will not b e observed. Ra the r , t he bubble will expand only slightly during crushing and achieve a volume approximately equal t o t h e to t a l inclusion volume. Therefore, if individual inclusions a r e not monitored during t h e crushing t e s t , and this is generally not possible with inclusions from epi thermal deposits because of t h e small s ize of t he inclusions and because of t h e randomness of t h e crushing procedure, the slight expansion of the vapor bubble t o fill t he inclusion will be missed. This probably explains why evidence of gases is almost never revealed during crushing t e s t s on inclusions from epi thermal deposits. Moreover, if an inclusion has an internal C 0 2 pressure of less than 10 bars a t roorn temperature , t h e c l a th ra t e compound does not form on cooling (Hedenquist and Henley, 1985). Because fluid inclusions trapped in the epi thermal environment generally have internal C O pressures below this value (Fig. 5.151, the presence of C 0 2 will not be revealed by recognition of the C 0 2 c la thra te .

The fourth technique for identifying C 0 2 , or o ther gases, in inclusions from t h e epi thermal environment i s microanalysis of t h e inclusion fluids. Virtually all reported occurrences of volatiles in inclusions from the epi thermal environment a r e a result of microanalytical techniques, and t h e majority of these a r e from mass spec t romet r i c analyses. The major problem with this approach is t h a t these analyses generally a r e bulk analyses of t he fluids released from a large number of inclusions, e i ther by crushing or thermal decrepitation. Because most samples from epi thermal deposits contain numerous growth zones and/or planes of inclusions representing fluids of potentially d i f ferent compositions trapped a t different t imes, t he resulting analyses probably a r e not representa t ive of any one fluid type, but ra ther mixtures of fluids of d i f ferent compositions and, perhaps, d i f ferent origins. Recent ly , however, Sommer et al. (1985) described a laser-decrepitation- capaci tance manometer technique which permits individual inclusions t o be sampled. Using this technique, these workers determined t h a t t h e C02!H20 ra t io of individual inclusions f rom t h e Topia, Mexlco epithermal precious-metal deposit was 0.0305 * 0.004.

INTERPRETATION OF FLUID INCLUSIONS FROM THE EPITHERMAL ENVIRONMENT

The purpose of most fluid-inclusion s tudies is t o obtain inforrnation on the temperatures , pressures (depths), and compositions of t he fluids responsible for mineral deposition, and t h e variation in these properties in t ime and space in individual hydrothermal systems. With currently available opt ica l equipment

Isobars at 25°C of Internal Pressures of

H20 - C02 Fluid Inclusions

Trapping Temperature (OC)

Figure 5.15. Calculated internal pressures (at 25O~) of H20-C02 inclusions with compos i- tions of 0.2 molal (solid lines) and 0.5 molal (dashed lines) C02 trapped at various epithermal P-T conditions. Also shown are the boiling curves for the fluids of these compositions, as well as the pure H20.

and heat inglf reezing s tages , i t is possible t o col lec t a large amount of highly accura t e d a t a f rom fluid inclusions in a relatively short period of t ime, and generally these d a t a consist of homogenization temperatures and melting temperatures of inclusion fluids. Mineralization conditions predicted f rom these da ta are , however, o f t en suspect, owing t o incorrect assumptions made during d a t a interpretation. In the relatively low-pressure epi thermal environment, fluid- inclusion homogenization t empera tu res require l i t t le or no pressure correct ion t o obtain t h e trapping temperature . Thus, homogenization temperatures provide a good approximation of t h e mineralization temperature and a r e not usually subject t o er rors in interpretation. The major e r ro r s in d a t a in terpre ta t ion a r e consequently re la ted t o composition and pressure determinations.

Generally, t h e inclusion composition is repor ted a s a salinity in t e rms of wt.-% NaCl equivalent and is obtained from t h e melting t empera tu re of i ce in the inclusion. The assumption t h a t is made in obtaining this salinity is t h a t t h e depression of the f reezing point is due solely t o dissolved solids, mainly chlorides of sodium, potassium, and calcium. However, a s

discussed above, some inclusions f rom the epi thermal environment may also contain smal l amounts of gases t h a t a r e not usually de tec t ab le by normal petrographic, microthermometr ic , or crushing studies, and these gases also contr ibute t o t h e f reezing point depression. Thus, in t h e pure H 0 - C 0 2 system, dissolved C 0 2 can result in lowering 0% the ice-melting t empera tu re t o -1.48OC, without formation of liquid CO or t h e C 0 2 h y d r a t e 5 314 H20) during cooing.

Hedenquist and Henley (1985) have shown t h a t this behavior may account fo r t h e apparently too high salinit ies determined from fluid-inclusion f reezing t empera tu res for many ac t ive geothermal systems. For example, inclusions f rom well BR25 a t Broadlands have a f reezing point of -0.6'~, which corresponds t o an NaCl equivalent salinity of 1.02 wt.%. Hedenquist and Henley (1985) have shown, however, t h a t 87% of t h e f reezing point lowering is due t o dissolved C O and t h e ac tual dissolved sa l t concentration is only 0.f; wt.%. The considerable e f f e c t t h a t this e r ro r in ca lcula ted chloride concentration has on pH and me ta l solubility e s t ima tes i s discussed by these workers.

Failure t o recognize small amounts of gases in fluid inclusions can also seriously a f f e c t t h e pressures and corresponding depths of format ion calculated f rom inclusion data. Most calculations of depths of

Temperature ("C)

Figure 5.16. Relationship between temperature and the depth at which boiling will commence for an H20-NaC1 solution ( 2 wt.-% NaC1) and H20-NaC1-C02 solutions containing 0.2 and 0.5 molal C02 and 0.5 molal NaC1. ( M d i f ied from Loucks, 1984.)

format ion from fluid-inclusion d a t a use the boiling point-depth relationship for H20-NaCI calcula ted by Haas (1971). Using these data , pressures calculated for epi thermal fluid inclusions in the t empera tu re range 200-300 '~ a r e on t h e order of a f ew t o several tens of bars. Depths obtained from these relatively low pressures a r e on t h e order of a f ew hundred me te r s t o a kilometer a t most for homogenization t empera tu res of 2 0 0 - 3 0 0 ~ ~ , a s shown by t h e temperature-depth curve for a 2 wt.-% NaCl solution on Figure 5.16. The addition of even small amounts of C 0 2 t o H20-NaCI, however, significantly raises the vapor pressure and, concomitantly, t h e calculated depth of formation. For example, addition of 0.5 molal C 0 2 t o a 0.5 molal (2.84 wt.%) NaCl solution a t 2 5 0 ' ~ raises the vapor pressure from 69 bars t o 108 bars, which increases t h e depth t o t h e boiling curve from 450 me te r s t o 1120 me te r s (Loucks, 1984). Thus, th is small amount of C O , which is in t h e range of C 0 2 concentrations generafly repor ted for ac t ive geothermal sys tems and fluid inclusions f rom epi thermal deposits (Fig. 5.141, increases the ca lcula ted depth of format ion a lmost 2 112 t imes a s compared t o depths ca lcula ted assuming only H20- NaCl.

APPLICATION OF FLUID INCLUSIONS IN EXPLORATION FOR EPITHERMAL

PRECIOUS-METAL DEPOSITS

Imaginative application of conceptual models of o re genesis is of paramount importance t o the contemporary exploration geologist. Over the las t decade, through a synthesis of detailed investigations of ac tua l deposits, experimental studies, fluid- inclusion research, and studies of ac t ive geothermal systems, a generalized model for t he format ion of epi thermal ore deposits has evolved. According t o some authors, boiling is an integral pa r t of th is model, and is thought t o be in t imate ly associated with mineralization in relatively shallow, near-surface deposits. Theoretical and exper imenta l studies (Drummond and Ohmoto, 1985) have shown t h a t under ce r t a in conditions, changes in the fluid chemistry a s a result of boiling could be a n e f f ec t ive mechanism for the precipitation of me ta l s from solution. In addition, t h e character is t ic low-pH al tera t ion assemblages commonly associated with these epi thermal sys tems a r e consistent with t h e "boiling off" and subsequent recondensation of ac idic gases, mostly C 0 2 and H2S, in t h e presence of cooler groundwater above the veins.

If boiling and mineralization a r e genetically re la ted , then, from an exploration geologist's viewpoint, fluid-inclusion evidence of boiling is a favorable character is t ic in t h e search for epi thermal deposits, and t h e pressures and corresponding depths ca lcula ted from inclusions trapped in a "boiling system1' define t h a t portion of t h e earth 's c rus t t h a t can host these deposits. The modification of depth e s t ima tes t o include t h e e f f ec t s of dissolved gases great ly expands t h e crus ta l depth range of "epithermal" precious-metal systems. However, t he ac tua l computation of depth es t imates becomes highly questionable without knowledge of gas con ten t s of

inclusions and the e f f e c t s of various gases on water- gas immiscibility.

In some deposits, i t can be demonstra ted unquestionably through geologic reconstructon t h a t mineralization occurred very near the surface (Vikre, 1985). For these surficial "hot-spring" deposits e i ther t he o re fluid had a very low gas content or, if t he o re fluid did contain dissolved gases and consequently began "boiling" a t some much deeper level, then near- surface deposition did not result from changes in fluid chemistry a s a result of boiling, but by some other mechanism such a s cooling, oxidation, or mixing.

A consideration of the e f f ec t s of dissolved gases on t h e depths a t which boiling will occur helps explain some apparent inconsistencies in the fluid-inclusion l i terature. If these deposits all formed very near t h e surface , a t or near t h e one-phaseltwo-phase (boiling) in ter face , then preservation of these deposits through geologic t i m e is unlikely. However, bonanza-type vein deposits and sediment-hosted disseminated deposits a r e not uncommon, and of ten occur in tectonically ac t ive a reas where uplift and erosion have taken place. Moreover, t he ver t ica l range over which mineralization occurs in individual deposits commonly extends t o depths considerably greater than those which could be achieved by simple H20-NaCI or H 2 0 hydrosta t ic boiling. For example, a t Guanajuato, Mexico, evidence for boiling can be found in samples col lec ted from 650 m below the present-day su r face (Buchanan, 1980). However, average fluid-inclusion t empera tu res of 2 3 0 ' ~ suggest a pressure of 28 bars (Haas, 19761, which conver ts t o on1 340 m of boiling Y . hydrosta t ic head ( avg. = 0.83 g/cm ). T h ~ s "deep-ore" problem is compounded if denudation s ince t h e t i m e of mineralization is accounted for, but is explicable if t h e e f f e c t s of dissolved gases a r e considered.

SUGGESTIONS FOR FUTURE FLUID-INCLUSION RESEARCH

Fluid inclusions f rom the epi thermal environment exhibit distinct petrographic and chemical character is t ics which serve t o distinguish them from those formed in deeper, higher t empera tu re systems. A t t h e present t ime, however, i t is no t possible t o r e l a t e these inclusion properties t o a specific type o r s t age of mineralization because a sufficient d a t a base does not exist , and the da ta t h a t a r e available a r e of ten inconclusive or contradictory. Many more deta i led fluid-inclusion studies of epi thermal sys tems (especially o r e deposits, but including barren sys tems a s well) will be required before these di f ferences can be reconciled. These studies should concen t ra t e on re la t ing observed fluid inclusions t o re la t ive s tages in a system's evolution, and variations in fluid-inclusion character is t ics within a system must be mapped in detail . Such careful, detailed studies of t h e temporal and spat ia l variations in fluid-inclusion character is t ics in t h e porphyry-copper deposits have contr ibuted great ly t o our understanding of the genesis of t hese deposits, and should prove equally valuable in understanding the epi thermal environment.

One a r e a in which additional work would be t imely is in defining the position of boiling in space

and t ime in epithermal sys tems and t h e relationship of boiling t o o r e deposition in mined o re deposits. I t is generally accepted t h a t boiling does occur in most epi thermal systems, but in those sys tems which have been studied in detail , t h e temporal and spatial relationship of boiling and o re deposition is highly variable. For example, a t Finlandia, Kamilli and Ohmoto (1977) found t h a t boiling only occurred a t t he location and paragenet ic s t age in which precious-metal deposition occurred. A t Sunnyside, boiling did not occur during gold mineralization, but was observed during the la ter quartz-rhodochrosite-fluorite s t age of mineralization (Casadevall and Ohmoto, 1977). Similarly, Rad tke e t al. (1980) repor t t h a t boiling a t Carlin was associated with l a t e acid leaching and hypogene oxidation, but did not occur during ear l ier quartz-pyrite-gold mineralization. A t Buckskin Mountain, Vikre (1985) found t h a t boiling was associated with precious-metal deposition a t shallow levels, but did not occur during precious-metal deposition a t deeper levels in t h e deposit. F rom these and other examples, i t is possible t h a t boiling may simply be a character is t ic of epi thermal sys tems in general, but may or may not be r e l a t ed t o precious- me ta l deposition. Thus, t he presence of boiling may be useful for identifying epi thermal sys tems bu t cannot, a t this t ime, be used t o predic t whether or not precious-metal deposition might have occurred.

A second, but related, a r e a in which m o r e work might be useful is t o document crys ta l habits and other character is t ics of various minerals fo rmed from boiling fluids, a s evidence of boiling may not be recorded by fluid inclusions. There probably a r e cer ta in conditions where fluid inclusions would not be trapped, or if formed, would no t record a l l fluid phases resulting f rom boiling. For example, t h e vapor phase resulting f rom near-surface boiling may no t b e t rapped in t h e minerals precipitating, but may simply escape through t h e system plumbing t o t h e surface. Hedenquist and Henley (1985) repor t t h a t ca l c i t e deposited in t h e casing of well 80 a t Wairakei during discharge of a high-temperature, two-phase mixture contained only liquid-rich inclusions. Also, t h e minerals precipitating a t near-surface conditions may be so f ine grained a s t o preclude observation of possibly sub-micron-sized inclusions.

Robertson (1968) has shown t h a t 91% of the inclusions trapped in sodium n i t r a t e crys ta ls grown in a boiling solution had liquid-to-vapor ra t ios suggesting t h a t t h e inclusions trapped only t h e liquid phase. Only 9% of the inclusions had variable liquid-to-vapor ra t ios indicating en t r apmen t of varying amounts of liquid and vapor. In a natura l sample in which 10% of the inclusions contain variable liquid-to-vapor ratios, these inclusions might easily be overlooked o r dismissed a s being inclusions which have necked or leaked, t h e conclusion being t h a t t h e inclusions were t rapped in a non-boiling system. Roedder (1984) descr ibes severa l o ther examples in which inclusions have preferentially t rapped only one of t w o coexisting fluid phases, result ing in inclusion character is t ics indicat ive of format ion in a single-phase fluid. Exper imenta l studies may help t o define those conditions under which only one type of inclusion may resul t during boiling, and should address the problem of inclusion

format ion in "slightly" boiling or ef fervesc ing systems, a s might occur when a fluid moves down a pressure gradient and slowly exsolves smal l amounts of vapor, a s compared t o inclusions fo rmed during violent boiling or "flashing," a s migh t occu r when a n overpressured sys tem is breached by development of fractures. The e f f e c t of gases should also be considered in t hese s tudies because t h e vapor phase produced during boiling is t h e non-wetting phase and may inhibit en t r apmen t a s inclusions.

REFERENCES

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Heald-Wetlaufer, P., Hayba, D. O., Foley, N. K., and Goss, J . A., 1983, Comparat ive anatomy of epi thermal precious- and base-metal d is t r ic ts hosted by volcanic rocks: U.S. Geological Survey, Open-File Repor t 83-710, 16 p.

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Khaibullin, I. K., and Borisov, N. M., 1966, Exper imenta l investigation of t he t he rma l proper t ies of aqueous and vapor solutions of sodium and potassium chlorides a t phase equilibrium: High Temperature , v. 4, p. 489-494.

Loucks, R. R., 1984, Zoning and o r e genesis a t Topia, Durango, Mexico: Unpublished Ph.D. thesis, Harvard, 416 p.

Malinin, S. D., and Kurovskaya, N. A., 1975, Solubility of C 0 2 in chloride solutions a t e levated t empera tu re s and C 0 2 pressures: Geokhimiya, no. 4, p. 547-550 (in Russian); t rans la ted in Geochemical International, v. 12, no. 2, p. 199-201.

Parkinson, W. J., and d e Nevers, N., 1969, Pa r t i a l molal volumes of carbon dioxide in wa te r solutions: Industrial and Engineering Chemis t ry Fundamenta ls , v. 8, p. 709-713.

Perkins, R. M., and Nieman, G. W., 1982, Epi thermal

R. J. BODNAR, T. J . REYNOLDS, & C. A. KUEHN 97

gold mineralization in t h e South Mountain volcanic dome, Summitvil le, Colorado; & Symposium of t h e Genesis of Rocky Mountain Ore Deposits: Changes with Time and Tectonics; Proceedings of Denver Region Exploration Geologists Society, p. 165-172.

Po t t e r , R. W., 11, and Brown, D. L., 1977, The volumetr ic proper t ies of a ueous sodium chloride 8 solutions f rom 0' t o 500 C a t pressures up t o 2000 bars based on a regression of available d a t a in t h e l i te ra ture : U.S. Geological Survey, Bulletin 1421-A, 36 p.

Quinn, E. L., and Jones , C. L., 1936, Carbon dioxide: American Chemical Society Monograph 72, Rheinhold, New York.

Radtke , A. S., Rye, R. O., and Dickson, F. W., 1980, Geology and s t ab l e isotope studies of t h e Car l in gold deposit , Nevada: Economic Geology, v. 75, no. 5, p. 641-672.

Ramboz, C., P ichavant , M., and Weisbrod, A,, 1982, Fluid immiscibil i ty in na tura l processes--Use and misuse of fluid inclusion data. 11. In terpre ta t ion of fluid inclusion d a t a in t e r m s of immiscibility: Chemical Geology, v. 37, p. 29-48.

Reed, M. H., and Spycher, N. F., 1985, Boiling, cooling, and oxidation in epi thermal systems: Numerical modeling approach; & Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemistry of Epi thermal Systems: Society of Economic Geologists, Reviews in Economic Geology, v. 2.

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solutions--an exper imenta l study: Unpublished M.S. thesis, University of Michigan, 53 p.

Roedder, E., 1984, Fluid Inclusions: Mineralogical Socie ty of America , Reviews in Mineralogy, v. 12, 644 p.

Sommer , M. A., 11, Yonover, R. N., Bourcier, W. L., and Gibson, E. K., 1985, Determinat ion of H 2 0 and C 0 2 concentra t ions in fluid inclusions in minerals using laser decrepi ta t ion and capaci tance manomete r analysis: Analytical Chemis t ry , v. 57, p. 449-453.

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Chapter 6 LIGHT STABLE-ISOTOPE SYSTEMATICS IN THE EPITHERMAL ENVIRONMENT

Cyrus W. Field and Richard H. Fifarek

INTRODUCTION

Stable-isotope geochemistry has made important contributions to t h e widely acknowledged "renaissance" in t h e e a r t h sciences for more than th ree decades. This s t a tus may be ascribed both t o theoret ica l and pract ica l considerations. First , t he isotopic species of an e lement may be f rac t ionated (partitioned unequally) between t w o o r more coexisting phases because of mass-dependent d i f ferences in their chemical and physical behaviors, and t h e amount of such fractionation normally varies inversely with temperature and independently of pressure. Accordingly, t he isotopic abundances of a n e lement may serve to define the mechanisms of formation, thermal environment, and provenance of rocks, minerals, and fluids. Second, t he analytical procedures now available render most geologic mater ia ls well suited for routine and rapid isotopic measurements. Some important milestones of t he 1930's and 1940's leading t o our present understanding include t h e discovery of deuterium and formulation of t h e theoretical basis for stable-isotope f rac t ionat ion by Harold C. Urey and colleagues a t t h e University of Chicago and the development of improved mass spect rometers by Alfred 0. Nier a t t h e University of Minnesota. The subsequent construction of laboratory facil i t ies elsewhere was commonly di rected by graduates and associates of these pioneers and thei r respective institutions.

As of today, t he l i tera ture re levant t o stable- isotope geochemistry is voluminous and f a r beyond the scope of this topical overview. Most investigations, apa r t from those concerned with theory o r laboratory experimentation, have been focused on one o r more of t he following objectives: (1) t h e conditions and mechanisms of rock or mineral formation; (2) t h e sources of magma, sediment, ore, .petroleum, or water and of metamorphic, ore-form~ng, or geothermal fluids; (3) t h e e f f ec t s of contamination re la ted t o magma-country rock, magma-water, o r water-rock reactions or t o natura l versus man-induced pollutants; and (4) t he geothermometry of rocks, ores, aqueous systems, and paleoclimates. Although many studies have deal t with the application of s t ab le isotopes t o mineral deposits and problems of o re genesis, relatively few have been concerned specifically and in deta i l with epithermal systems. This meager d a t a base, given the current overriding in teres t in deposits of t h e precious metals, undoubtedly will b e enlarged both in detail and scope during the forthcoming decade. Our emphasis in this review is confined exclusively t o t h e s table isotopes of carbon, hydrogen,

oxygen, and sulfur. Prior t o our discussion of the d a t a for epi thermal deposits, we provide a background synopsis of t he conventions of da t a presentation, fractionations t h a t may accompany equilibrium isotope-exchange react ions between various fluid and mineral phases, and t h e distributions of t h e s table isotopes in major geologic environments. Those who require more precise elaboration should consult recent textbooks by Faure (1977) and Hoefs (1980) and reviews by Friedman and OfNeil (19771, OfNeil (19771, Ohmoto and Rye (19791, and Taylor (19791, and appropr ia te references c i t ed therein.

CONVENTIONS, SYSTEMATICS, AND RATIONALE

According t o Lange's Handbook of Chemist ry (Dean, 1979), t he re la t ive abundances of t he s t ab le isotopes of hydrogen, carbon, oxygen, and sulfur a r e a s follows:

Although these abundances a r e not known t o t h e accuracy implied, this uncer ta in ty is not important because a precision of 0.02 percent or be t t e r is obtained routinely f rom analyses based on comparat ive

measurements of t h e heavy-to-light isotopic ra t io of an e l emen t in a sample with respect t o t h a t in a standard. The analyses a r e performed on a gas-source mass spec t romete r , and they require t h a t t h e isotopic e lement measured in both t h e sample and s tandard be converted quant i ta t ively t o a gas (carbon and oxygen t o C 0 2 , hydrogen t o H2, and sulfur t o SO ). Depending on t h e e l emen t under consideration, t i e instrumental record provides a com ris n of sample and s tandard ra t ios in t e rms of D/H, "CI'~C, 180/160, and 3 4 ~ / 3 2 ~ . Differences in the ra t io of t he sample (R ) with respect t o t h a t of t h e standard (Rs) a r e normafiy expressed a s deviations from the s tandard by del (6) values in pa r t s per thousand (permil, or O/oo), a s given by t h e equation

Posit ive or negative del values signify the permil enr ichment r d letion respectively, of t h e heavy isotope (D, lgC, '$, or in the sample re la t ive to t h e standard. Conventional standards a r e a belemnite from the Cretaceous Peedee Formation (PDB) for carbon, Standard Mean Ocean Water (SMOW) for hydrogen and oxygen, and t ro i l i te from the Canyon Diablo me teo r i t e (CD) for sulfur. All standards a r e 0 permil, by definition (eq. 1).

Fractionation

Isotopes of an e lement may be f rac t ionated during equilibrium and unidirectional reactions t h a t accompany chemical and physical processes of e i ther inorganic or biogenic origin. The fractionation e f f e c t s a r e generally d i rec t ly proportional t o mass d i f ferences between t h e isotopic species and inversely proportional t o temperature , provided other parameters such a s valence and bond s t rength a r e the same. Equilibrium isotope-exchange react ions a r e probably the most important , and many of the fractionations f i rs t predicted from theory have been corroborated by subsequent laboratory studies and analyses of natura l minerals, liquids, and gases. Fractionations t h a t accompany unidirectional reactions a re more difficult t o evaluate. They evolve from kinetic e f f ec t s and result in t h e reaction products being depleted in the heavy isotope.

The isotopic fractionation f ac to r ( a ) between t w o compounds, A and 9, is equal t o the quotient of t he heavy-to-light isotope ratios in the compounds. Moreover, under conditions of equilibrium, the fractionation f ac to r and ra t io quotient a r e re la ted to t h e equilibrium constant (K) of isotope exchange a s follows,

where n is t he maximum number of exchangeable isotopes in any of t h e compounds. However, this generalization with respect t o K does not hold t rue for

compounds of hydrogen or fo r those containing t w o or more isotopes of an e lement t h a t do not occupy equivalent molecular positions. Because instrumental measurements of t he stable-isotope abundances a re normally repor ted a s del (6) values in permil ra ther than a s ra t ios (R), o ther identit ies must be used t o equa te these deviations f rom t h e standard (6O/00) to t h e f rac t ionat ion f ac to r (a). The interrelationships given by

a r e derived from equations ( I ) and (2). Isotopic e f f ec t s deduced from exper imenta l and

theoret ica l investigations a r e usually reported in t e rms of t h e f rac t ionat ion factor . The a values a r e close t o unity and commonly range from I * O.Ox t o 1 i 0 . 0 0 ~ . In contras t , isotopic e f f ec t s obtained from applied s tudies of minerals and rocks a r e generally reported a s de l t a (A) values, which is t h e difference of measured del values, in permil, between two phases. Delta values may be readily determined from inspection of t h e analyt ica l d a t a (given a s 6 '100 values) and a r e similar t o values derived from the logarithmic t ransformat ion of their associated fractionation factors , a s provided by the useful approximation

For most pract ica l purposes, such a s geothermometry, t h e de l t a value is essentially identical t o t h e more precise, but less easily calculated, 1000 In a value. Errors introduced by this approximation (eq. 4) become significant, relative to analytical uncertainties (* 0.2 '100 or more), only a s the de l t a values exceed 10 permil and a s thei r component del values become unsymmetrically distributed with respect t o the reference standard (0 '100).

I l lustrative of t h e foregoing considerations, the equation

repr e n t s t equilibrium isotope-exchange reaction for "S and "S between sulfur dioxide and hydrogen sulfide gases. I t follows, from equations (I) , (21, and (3), t h a t

and f r a n e q u a t i o n ( 4 ) t h a t

The equilibrium fract ionat ion f ac to r s for this react ion a r e 1.0339 and 1.0074 a t 1 0 0 ~ ~ and 500°c, respectively, according t o t h e equation for SO2-H S equilibration given by Ohmoto and Rye (1979, p. 5 1 6 . Equivalent values of 1000 In ci(and A s e e eq. (7)) a r e b 33.3 and 7.4 permil a t 1 0 0 ~ ~ and 500 C, respectively. Thus, under equilibrium conditions, SO2 is enriched in the heavy 3 4 ~ isotope by 33.3 and 7.4 permil re la t ive t o coexisting H2S at these two t empera equation (5) been wri t ten in t e r m s of the isotopic species of SO2 and HZS, t h e associated f rac t ionat ion f ac to r s would have Increased t o 1.0678 and 1.0148 a t 1 0 0 ~ ~ and 500°C, respectively. Thus, a s previously s ta ted , fractionation e f f ec t s r e l a t e inversely t o t empera tu re and directly t o d i f ferences in mass between t h e isotopic species involved. However, small f rac t ionat ions between coexisting phases of a mineral group formed a t similar t empera tu s, such a s variations in 6 3 4 ~ among sulfides or 6 "0 among silicates, r e l a t e t o differing s t rengths of the meta l - sulfur or metal-oxygen bonds (Bachinski, 1969; O'Neil, 1977). The heavy isotope is preferentially concentra ted (enriched) in t h a t phase having t h e s t rongest bond.

Equilibrium React ions

The t empera tu re dependency of stable-isotope f rac t ionat ion provides geologists with a potentially large number of geothermometers. However, successful application to t h e study of rocks and minerals requires t h a t t h e isotopic e lement be geologically common, and t h a t t h e fractionation f ac to r s for t he isotopes between two or more phases be known over t h e appropr ia te range of temperatures. Although the most reliable equilibrium fractionation f ac to r s a r e obtained f rom laboratory investigations of isotope exchange, experimental problems involving disequilibrium and metastabili ty, particularly a t low temperatures , commonly necess i ta te t h a t they be e s t ima ted indirectly by less cer ta in methods. These e s t ima tes a r e derived e i ther from theory, using s ta t i s t ica l mechanics and various spect ra l and thermodynamic data , or from empirical consideration of measured fractionation trends in natural mineral assemblages. Fractionation equations for isotopic react ion pairs of many geologically important compounds of carbon, hydrogen, oxygen, and sulfur a r e presented in Tables 6.1 through 6.5. The general form of these equations

indicates t h a t t h e f rac t ionat ion e f f ec t , given by 1000 In a. varies curvilinearlv with the inverse sauare of

in Figures 6.1 through 6.5. Posit ive o r negat ive values of 1000 In ci ( ~ A v a l u e s ) denote permil enrichment o r depletion, respectively, of t h e heavy isotope in t h e first-named member of t h e pair. For purposes of illustration, some of the f rac t ionat ion curves have been extrapola ted t o t empera tu res beyond t h a t of t he exper imenta l d a t a and somet imes beyond t h a t of t he stabil i ty range fo r a t leas t one member of t he react ion pair.

Fract ionat ion equations for react ion pairs of isotopic carbon compounds a r e listed in Table 6.1 and t h e resul tant curves over t h e t empera tu re range from oOC t o 600°c a r e shown in Figure 6.1. They a r e based ent i re ly on the sources of d a t a compiled by Friedman and O'Neil (19771, although Ohmoto and R y e (1979) provide more complex equations for these and other carbon compounds. As previously noted, t he heavy isotope is preferentiallv concentra ted in t h a t member of t h e react ion pair t o which the e lement is most strongly bonded. Equilibrium values of 1000 (Fig. 6.1) f o r t h e carbon react ion pairs show t h a t @ tends t o b e enriched in t h e more condensed and(or) oxidized member. The l a t t e r redox e f f ec t , also common t o isotopic compounds of sulfur, i s confirmed by d a t a from sedimentary, tamorphic, and geothermal environments wherein T'C is found t o be concentra ted in carbonates re la t ive t o graphite or organic carbon, and in C 0 2 re la t ive t o CH4.

The fractionation equations for isotopic compounds of hydrogen a r e given in Table 6.2 and t h e portrayal of these e f f e c t s with temperature in Figure 6.2 a r e largely f rom t h e compilation of Friedman and OINeil (1977). However, more recent exper imenta l d a t a for some mineral-H20 equilibria include those for

CARBON

-- t empera tu re (in OK). doe f f i c i en t s a and b r ip re sen t 0 100 ZOO 300 400 500 600 t h e slope and in tercept , respectively, of these linear TEMPERATURE OC curves. Variations in the fractionation e f f e c t (1000 Ins) over t h e t empera tu re interval 0 t o 6 0 0 ' ~ for F i g u r e 6.1. F r a c t i o n a t i o n c u r v e s f o r isotopic many of these react ion pairs a r e graphically i l lustrated compounds o f carbon.

Table 6.1 Fractionation equations for isotopic compounds of carbon

Carbon Pair 1000 lna T (OC) Range Reference

1 C02-CH4 not given 0-700 Friedman and O'Neil (1977, Fig. 29)

2 gph-CH4 not given 0-700 Friedman and O'Neil (1977, Fig. 32)

3 cal-gph '1.74(lo6/T2) + 5.22 0-700 F r i e d m a n a n d O f N e i l ( 1 9 7 7 , F i g . 2 9 )

4 C02-gph not given 0-700 Friedman and O'Neil (1977, Fig. 30)

5 cal-C02 2.988(106/T2) - 7.666(103/T) 0-700 Friedman and O'Neil (1977, Fig. 31) + 2.461

6 HCO;-CO~ 9.552(103/~) - 24.10 5-125 Friedman and O'Neil (1977, Fig. 27)

7 CO;-CO~ not given (-eq. 6) 0-100 Friedman and O'Neil (1977, Fig. 28)

8 dol-cal 0.18(lo6/T2) + 0.17 100-650 Friedman and O'Neil (1977, Fig. 33)

Comments

1. Although Friedman and O'Neil (1977) and sources cited therein do not provide fractionation equations for some of the carbon pairs listed above, their graphical portrayal of fractionation trends for these pairs are the basis of those curves illustrated in Figure 6.1 and our approximation of the calcite-graphite equation (eq. 3) above.

2. Ohmoto and Rye (1979, p. 551, Table 10-3) give more complex polynomial equations for the equilibrium fractionation factors of these and other isotopic compounds of carbon.

3. Fractionation equations and isotopic trends for the carbon pairs have been calculated largely from empirical and theoretical considerations, and not from experimental determinations, over the temperature ranges listed (see Friedman and O'Neil, 1977, and Ohmoto and Rye, 1979).

4. Mineral abbreviations and compound states are as follows: cal, calcite; CH4, gaseous; COZY gaseous; CO;, aqueous; dol, dolomite; gph, graphite; HCO;, aqueous.

epidote (Graham e t al., 19801, goethi te (Yapp and Pedley, 19851, and hornblende (Graham et al., 1984). Less cer ta in and(or) complete a r e t h e d a t a fo r gibbsite and kaolinite (Taylor, 1979) and for montmoril lonite (Savin and Epstein, 1970a; O'Neil and Kharaka, 19761, and the re a r e virtually none for chlorite. Equations and curves a r e not given fo r equilibrium react ions of H2-H20(v1, H2-CH4, and H2S-H20. However, the i r fractlonatlons a r e enormous, a s summarized by Friedman and OtNeil (19771, with 1000 In values ranging from -900 and less t o -400 and less within the temperature interval f rom OOC t o ~ O O ~ C , respectively. Presumably react ions such a s those with H a r e no t important in most geologic environments (T%ylor, 1974a; 1979). Values of 1000 In a r e largely negative, fo r t he equations (Table 6.2) a s writ ten and thei r curves (Fig. 6.21, and they demonstra te t h a t D is enriched in H 2 0 relative t o coexisting gas or mineral phases. The parallelism of f rac t ionat ion curves for t h e various mica-water sys tems (Fig. 6.2, eqs. 5, 6, and 7) is noteworthy. According t o Suzuoki and Epstein

(1976), t he f rac t ionat ion of D between t h e micas and water i s not only a function of temperature , but also r e l a t e s t o t h e molar f rac t ion of cations (Al, Fe , and Mg) in six-fold coordination. From their general equation fo r mica-water f rac t ionat ion given by

1000 l n a

where X is t h e molar f rac t ion of each cation, i t may be inferred t h a t values of 1000 l n a a t constant t empera tu re become progressively more depleted (negative) in D with t h e compositional succession from Al-rich through Mg-rich t o Fe-rich micas. In t e rms of theoret ica l mineral assemblages, Al-pure "muscovite" would be enriched in D by 6 permil re la t ive t o coexisting Mg-pure "phlogopite", and by 70 permil re la t ive t o coexisting Fe-pure Itannite"; regardless of temperature . We have se l ec t ed compositionally reasonable molar f rac t ions of these cations

C . W. FIELD & R. H. FIFAREK 103

(XAI:XM<XFe) for t h e fractionation equations and curves 6 and 7) i l lustrative of phlogopite (0.10:0.80:0.10) and bioti te (0.10:0.45:0.45) in Table 6.2 and Figure 6.2. These and(or) possibly other compositional variations may similarly a f f e c t the distributions of hydrogen isotopes in chlorites and smectites.

Fractionation equations and curves for equilibrium isotope-exchange reactions involving oxygen compounds a r e presented in Table 6.3 and Figure 6.3, respectively. Although many of these reactions a r e from sources in Friedman and O'Neil (19771, they include d a t a for reactions of qua r t z and the feldspars f rom Matsuhisa e t al. (19791, kaolinite from Kulla and Anderson (1978), and bioti te and chlor i te from Bottinga and Javoy (1973) and Javoy (1977) based on empirical and theoret ica l considerations. Limited da ta f rom Sheppard e t al. (19691, Savin and Epstein (1970a), and Taylor (1979) suggest t h a t t h e equation for montmoril lonite-water fractionation may be grossly analagous t o t p ~ t for bar i te water. Note t h a t t h e fractionations of 0 and given by 1000 Ina values for t h e various mineral-water sys tems (Fig. 6.3) a r e substantially less than those for deuterium. With t h e exceptions of bioti te, chlor ' e and magnetite, most minerals concen t ra t e l i d re la t ive t o coexisting water, which is a reversal of t h e isotopic e f f e c t tor deuterium. The order of enr ichment among t h e sil icates is largely t h e rev rse of Bowen's reaction series, with qua r t z enriched in "0 rela t ive t o most o ther common rock-forming minerals. Fractionation equations (Table 6.3) for a lb i te and potassium feldspar a r e identical, a s a r e thei r respective 1000 l n a values (Fig. 6.3) fo r a given

-1u 0 100 200 300 400 500 600

TEMPERATURE OC

Figure 6.3. Fractionation curves for isotopic compounds of oxygen.

t empera tu re In contras t , anor th i t e i s slightly depleted in ''0 rela t ive t o t h e alkali feldspars a t a l l temperatures , which r e f l ec t s a progressive weakening of t h e Si -0 bond with increasing substi tution of alurninum for silicon and sodium for calcium in the solid solution sequence of t h e plagioclase feldspars.

Equations and curves pertaining t o equilibrium fractionations between isotopic compounds of sulfur a r e listed in Table 6.4 and a r e graphically i l lustrated in Figure 6.4. The equation fo r sulfate-H S fractionation (eq. I ) is from Ohmoto and Lasaga (19$2), whereas a l l o thers a r e those provided by Ohmoto and Rye (1979). Large f rac t ionat ions t h a t cha rac te r i ze the sulfate-H S and SO2-H2S react ions a r e re la ted primarily t o t i e d i f ferent s t a t e s of sulfur xidation among the members of these pairs, with 38S being preferentially enriched in t h e more oxidized member of t he pair (Fig. 6.4, inse t curves 1 and 2). They a r e analagous t o t h e large isotopic e f f e c t previously described for carbon t h a t a t t ends t h e C O CH4 redox reaction (Fig. 6.1, curve I). The smal ler 6 0 0 lna values for sulfide- H2S react ions a s contras ted with those for si l icate- H 2 0 react ions is par t ly a t t r ibutable t o t h e s aller percentage mass d i f ference betwe '% and 'S a s compared t o t h a t between 1 8 0 and "0. I t is apparent from the signs of t h e 1000 lna values fo r various sulfide-H2S react ions (Fig. 6.4) t h a t molybdenite, p y r j p , and sphaler i te a r e variably but weakly enriched in S re la t ive t o coexisting H S, whereas galena and the other sul ldes a r e depfeted. The re la t ive 5 1 enr ichment of S between di f ferent sulfide minerals

TEMPERATURE oc (mo>py>sl, etc.) is governed by the re la t ive s t rengths of t h e metal-sulfur bonds, a s f i rs t predicted from

Figure 6.2. Fractionation curves for isotopic theory by Sakai (1968) and Bachinski (1969).

compounds of hydrogen.

104 CHAPTER 6

Tab le 6.2 F r a c t i o n a t i o n e q u a t i o n s f o r i s o t o p i c compounds of hydrogen

Hydrogen P a i r 1000 lncl T (OC) Range Refe rence

-24.844(lo6/T2) + 76.248(103/T) 0-50 Friedman and O'Neil (1977, F igs . 34 and 35)

2 CH4-H20(1) n o t g iven

3 CH4-H20(,) n o t g i v e n

0-350 Friedman and O'Neil (1977, F igs . 34, 35, and 36)

0-700 Friedman and O'Neil (1977, Fig. 36)

4 hbd-H20 -23.1 350-805 Graham e t a l . (1984)

5 mus-H20 -22.1(lo6/T2) + 19.1 450-850 Suzuoki and E p s t e i n (1976)

6 phl-H20 -22.4(106/T2) + 28.2 + C , where* 450-850 Suzuoki and E p s t e i n (1976) C = -9.8

7 bio-H20 -22.4(106/T2) + 28.2 + C, where* 450-850 Suzuoki and E p s t e i n (1976) C = -32.2

<300 Graham e t a l . (1980) 300-650

9 gib-H20 -16.1 a t 20 Taylor (1979)

10 kal-H20 n o t g i v e n 20-300 Taylor (1979, F i g . 6.2)

11 mon-H20 -64 and -43.9 15 and 200 Savin and E p s t e i n (1970a) and O'Neil and Kharaka (1976)

12 goe-H20 -104 and -103 0 and 100 Yapp and Ped ley (1985)

Comments

1. F r a c t i o n a t i o n cu rves i l l u s t r a t e d i n F i g u r e 6.2 f o r hydrogen p a i r s l a c k i n g f r a c t i o n a t i o n e q u a t i o n s a r e based on, o r determined from, t h e g r a p h i c a l p o r t r a y a l of t h e s e i s o t o p i c t r e n d s g i v e n by Friedman and O'Neil (1977) and T a y l o r (1979).

2. Mine ra l a b b r e v i a t i o n s and compound s t a t e s a r e a s f o l l o w s : b i o , b i o t i t e ; CH4, gaseous ; e p i , e p i d o t e ; g i b , g i b b s i t e ; goe , g o e t h i t e ; H20, l i q u i d u n l e s s o the rwise s p e c i f i e d ; hbd, ho rnb lende ; k a l , k a o l i n i t e ; mon, m o n t m o r i l l o n i t e ; mus, muscovi te ; and p h l , p h l o g o p i t e .

* Where C i s c a l c u l a t e d u s i n g t h e g e n e r a l e q u a t i o n : C = 2XA1 - 4XMg - 68XFe. See t e x t o r Suzuoki and E p s t e i n (1976) f o r an e l a b o r a t i o n .

Applications

Geologists who a r e unfamiliar with t h e innuendos of stable-isotope geochemistry may question our apparently excessive concern with exper imenta l and(or) theoret ica l sys tems defined by mineral-H20 and mineral-H2S reactions. Jus t i f ica t ion for this approach comes from t h e f a c t t h a t t h e results of such laboratory investigations offer guides and const ra in ts t o t h e in terpre ta t ion of stable-isotope data , and ideally, t hey provide the means by which some re la t ive

genet ic pa ramete r s of rock and mineral sys tems may be quantified. Two prominent applications will now be described in order t o s t ress t h e importance of laboratory research. These and other uses will be implicit t o subsequent discussions and in terpre ta t ions of t he stable-isotope data.

Geothermometry--Although the t empera tu re dependency of f rac t ionat ion serves a s t h e ra t ionale for t h e application of s table isotopes t o geothermometry , t h e acquisition of reliable d a t a from geologic mater ia ls requires t h a t severa l prerequisites be

C. W. FIELD & R. H. FIFAREK 105

Tab le 6.3 F r a c t i o n a t i o n e q u a t i o n s f o r i s o t o p i c compounds of oxygen

Oxygen P a i r 1000 l n a T (OC) Range Refe rence

C02(g)-H20(1) - 0 . 0 2 1 ( 1 0 ~ / ~ ~ ) + 17 .994(103/T) 0-100 Friedman and O'Neil (1977, - 19.97 F igs . 5 and 6 )

100-575 Friedman and O'Nei l (1977, Fig. 14)

250-500 Matsuhisa e t a l . (1979) 500-800 Matsuhisa e t a l . (1979)


<180-350 Friedman and O'Neil (1977, Fig. 15)

6 C02(g)-cal -1.803(lo6/T2) + 1 0 . 6 1 1 ( 1 0 ~ / ~ ) 0-750 Friedman and O'Neil (1977, - 2.780 Fig. 12)

9 mon-H20 eq . 5



15-300 Sheppard e t a l . (1969) , Savin and E p s t e i n (1970a) , and Taylor (1979)


11 kal-H20 2.05(lo6/T2) - 3.85 <150-300 K u l l a and Anderson (1978)

0-300 Friedman and O'Neil (1977, Fig . 8 )

1 3 bio-H20 0 .41( lo6/T2) - 3.10 ( ? ) 500-800 Bo t t inga and Javoy (1973) and Javoy (1977)

14 chl-H20 -1.34(lo6/T2) - 2.07 ( ? ) 500-800 Bo t t inga and Javoy (1973) and Javoy (1977)

500-800 Bo t t inga and Javoy (1973) and Javoy (1977)

Comment

1 . Mine ra l a b b r e v i a t i o n s and compound s t a t e s a r e a s f o l l o w s : Ab, a l b i t e ; An, a n o r t h i t e ; anh , a n h y d r i t e ; b a r , b a r i t e ; b i o , b i o t i t e ; c a l , c a l c i t e ; c h l , c h l o r i t e ; C02, gaseous ; H20, l i q u i d u n l e s s o the rwise s p e c i f i e d ; k a l , k a o l i n i t e ; Kf , K-fe ldspar ; mag, m a g n e t i t e ; mon, mon tmor i l lon i t e ; mus, muscovi te ; and q t z , q u a r t z .

fulfilled. These include not only a knowledge of fo rmed contemporaneously and in isotopic equilibrium, variations in t h e fractionation f ac to r with t empera tu re and t h a t this equilibrium was subsequently preserved. between two or more mineral phases, a s determined These fundamental t ene t s notw' hstanding, because preferably frorn experiment, but also t h a t t h e minerals t he f rac t ionat ion of D and " 0 in s i l ica tes i s

Figure 6.4. Fractionation curves o compounds of sulfur. 5

0 100 200 300 400 500 600 TEMPERATURE OC

Table 6.4 Fractionation equations for isotopic compounds of sulfur

Sulfur Pair 1000 lncl T (OC) Range Reference

for isotopic

6.463(10~/~~) + 0.56 -200 to >600

4.70(106/T2) - 0.5 350-1050

0.45(106/T2) uncertain

0.40(106/T2) 200-700

o.~o(~o~/T~) 50-705

-0.05(10~/~~) 200-600

-0.16(106/T2) 200-400

-0.25(10~/~~) uncertain

-0.63(106/T2) 50-700

-0.70(106/T2) uncertain

-0.75(106/T2) uncertain

eq. 11 uncertain

-0.80(10~/~~) uncertain

Ohmoto and Lasaga (1982)

Ohmoto and Rye (1979)












Comment

1. Mineral abbreviations and compound states are as follows: ag, argentite; bn, bornite; cb, cinnabar; cc, chalcocite; cp, chalcopyrite; gn, galena; H2S, gaseous; mo, molybdenite; pyl pyrite; SO, native sulfur; sb, stibnite; sl, sphalerite; SO2, gaseous; SO;, either aqueous or solid sulfate.


referenced t o s i l ica te-H20 reactions and t h a t of 3 4 ~ in sulfides t o sulfide-H2S reactions, i t follows t h a t more geologically useful equat ions fo r t h e fractionation of these isotopes between mineral phases, such a s between sil icate-sil icate and sulfide-sulfide pairs, may

derived f rom algebraic identities. For example, the "0 fract ionat ion equations given for t he quar tz-H20 and bar i te-H20 equilibrations (Table 6.3, eqs. (3) and (5)) a r e likely t o b e of in teres t only t o some geothermalists, whereas thei r algebraic d i f ference

1000 l n a qtz-H20 - 1000 Ina bar-H20

yields the f rac t ionat ion equation for t he quartz-barite reaction, which is an assemblage of in r e s t t o many explorationists. By having analyses of $0 performed on the members of this pair, and applying the approximation given by equation (4)

'OoO lna q t z -ba r " A q t z - b a r

= alsoqtz OIOO - ~ 1 8 % ~ ~ O/oo (11)

the depositional t empera tu re (in OK) of the quartz- bar i te assemblage is readily calculated by substi tution and solving fo r T in equation (10).

Listed in Table 6.5 a r e the equilibrium fractionation equations, and their analogues rearranged in t e r m s of T, for isotopic mineral pairs of oxygen and sulfur in potentially useful assemblages. Variations of thei r respective 1000 l n a (=A) values over the t empera tu re range from O°C t o 6 0 0 ' ~ a r e i l lustrated in Figure 6.5. The equations l isted have e i ther been calcula ted f rom those in Tables 6.1-6.4 by t h e method described (eq. 101, o r they have been taken from sources previously noted such a s Friedman and OfNeil (19771, Matsuhisa e t al. (1979), Ohmoto and Rye (19791, e tc . Many other equations for mineral-mineral reactions can be derived from those provided in Tables 6.1-6.4, but those given in Table 6.5 a r e res t r ic ted t o mineral pairs commonly found in hydrothermal deposits.

The use of t hese and other f rac t ionat ion equations in isotope geothermometry requires fur ther elaboration. Because of problems with disequilibrium t h a t a r e especially prevalent a t low reaction temperatures , not a l l of t h e equations in Tables 6.1-6.5 a r e equally reliable fo r isotope geothermometry. For example, t h e quar tz-H20 reaction has been repeatedly investigated (see Fr iedman and O'Neil, 1977; O'Neil, 1977; Matsuhisa e t al., 19791, because t h e exchange of oxygen isotopes between qua r t z and water is particularly slow, and probably most o ther sys tems will be reexamined eventually. For similar reasons, isotopic equilibrium is unlikely t o be a t ta ined with sulfate-sulfide react ions below 350°c, although i t may be enhanced by conditions such a s long residence t i m e and(or) low r a t e of cooling, low pH, and high sulfur content of t he fluids (Ohmoto and Rye, 1979; Ohmoto and Lasaga, 1982). Pyrite-chalcopyrite pairs commonly give unreasonable isotopic temperatures

TEMPERATURE OC

Figure 6.5. Fractionation curves for isotopic mineral pairs of oxygen and sulfur.

indicative of disequilibrium (Field and Gustafson, 19761, whereas sphalerite-galena pairs generally do not (Ohmoto and Rye, 1979). Furthermore, t h e carbon- bearing minera ls a r e not only few in number, but t h e f l a t slope of t h e dolomite-calcite fractionation curve with respect t o t empera tu re renders this geologically common mineral pair of l i t t l e value a s a geothermometer ; in contras t t o t h e s teeper slope of the calcite-graphite fractionation curve (Fig. 6.1, curves 8 and 3). Application of hydrogen isotopes t o geothermometry is severely l imited because of uncer ta in t ies caused by compositional variations of t h e minerals, post-depositional re-equilibration, and for some react ions t h e insensitivity of fractionation t o changes in t empera tu re (Fig. 6.2). The 1 8 0 compositions of t h e sil icate minerals a r e subject t o re- equilibration by aqueous fluids, and this post- depositional e f f e c t is larges t for t h e feldspars and bioti te, and l eas t for quar tz and magnet i te (Brigham, 1984, and references c i ted therein). Finally, t h e geologist must exerc ise ca re in the selection of samples fo r analysis. Although minerals t h a t exhibit t ex tu ra l evidence of noncontemporaneity a r e unlikely t o be in isotopic equilibrium, t h e converse i s not always true.

Provenance--Geochemical investigations commonly have a s thei r objective t h e identification of t h e source or provenance of fluids and(or) minerals by means of chemical tracers. Stable isotopes have been successfully applied t o many of these endeavors. In particular, t h e isotopes of hydrogen and oxygen a r e used routinely t o identify the source of aqueous fluids,

108 CHAPTER 6

Tab le 6.5 F r a c t i o n a t i o n and t empera tu re e q u a t i o n s f o r i s o t o p i c m i n e r a l p a i r s of oxygen and s u l f u r

Mine ra l P a i r 1000 l n a A TOK

1 q tz-mag 4.81 ( lo6 /T2) + 0.39 (250-500°C) 2 .19 ( lo3 ) / (n - 0 . 3 9 ) l l 2 3 .52( lo6/T2) + 2.56 (500-80o0c) 1 . 88 (103) / (n - 2 . 5 6 ) l l 2

2 q t z -ka l 1.29(106/T2) + 0.54 (250-500°c) l . 1 3 ( l o 3 ) / ( ~ - 0 . 5 4 ) l / ~

3 q tz-bar 0.34(lO6/T2) + 3.48 (250-500°C) 0 . 5 8 ( 1 0 3 > / ( ~ - 3 . 4 ~ ) ~ ~ ~ 0.95(lO6/T2) + 5.65 (500-80o0c) -0 .97(103) / (n - 5 . 6 5 ) l I 2

4 qtz-mus 0 .96( lo6/T2) + 0.58 (250-50o0c) 0 .98 ( lo3 ) / (n - 0 . 5 8 ) l / ~ -0.33(lO6/T2) + 2.75 (500-650'~) - o . 5 7 ( l o 3 ) / ( ~ - 2 . 7 5 ) l l 2

6 anh-qtz 0.54(106/T2) + 0.41 (250-500°C) o . 7 3 ( i o 3 > / ( ~ - 0 . 4 1 ) ~ ~ ~ 1.83(lO6/T2) - 1.76 (500-800'~) 1 . 3 5 ( l o 3 ) / ( n + 1 . 7 6 ) l l 2

7 q t z - c a l 0 .56( lo6/T2) - 0.42 (250-500°C) 0 . 7 4 ( i o 3 ) / ( ~ + 0 . 4 2 ) l / ~ -0.73(lO6/T2) + 1.75 (500-800°C) - 0 . 8 5 ( l o 3 ) / ( ~ - 1 . 7 5 ) ~ ~ ~

10 PY-ag 1 . 2 0 ( 1 0 ~ / ~ ~ ) ( u n c e r t a i n ) l . l o ( l ~ ~ ) / ( ~ ) ~ / ~

1 2 s l -gn 0.73(lO6/T2) (50-70o0c) 0 . 8 5 ( 1 0 ~ ) / ( ~ ) ~ / ~

Comments

1. These e q u a t i o n s a r e d e r i v e d from t h o s e l i s t e d i n Tab les 6.3 and 6.4 by methods d e s c r i b e d i n t h e t e x t .

2. Mineral and compound a b b r e v i a t i o n s a r e a s f o l l o w s : Ab, a l b i t e ; a g , a r g e n t i t e ; anh, a n h y d r i t e ; b a r , b a r i t e ; c a l , c a l c i t e ; c p , c h a l c o p y r i t e ; gn, g a l e n a ; k a l , k a o l i n i t e , Kf , K-fe ldspar ; mag2 m a g n e t i t e ; mus, muscovi te ; py, p y r i t e ; q t z , q u a r t z ; s l , s p h a l e r i t e ; and SO4 e i t h e r aqueous o r s o l i d s u l f a t e .

whether o r not now present, in many rock- and ore- forming environments. The viability of this method s t e m s from the f a c t t h a t t h e isotopic domains of magmat ic , meteor ic , and ocean wa te r s a r e generally distinct, and t h a t the composition of fluids (hydrothermal, metamorphic, etc.) may be calcula ted from t h e analytical d a t a for one or more minerals using the appropriate reaction equations (Tables 6.2 and 6.3). For example, given t h e AD and 6180 values fo r a sample of hydrothermal ser ic i te , and assuming equilibrium and a temperature of format ion based on independent cr i ter ia (fluid inclusions, mineral

assemblages, sulfur isotopes, etc.), t h e compositions of t h e fluid may be determined by substi tution of the d a t a and assumed t empera tu re (in OK) in t h e muscovite-H20 react ion equations ( able 6 .2 , eq. (5 ) ;

I 6 Table 6.3, eq. (8)). Similarly, t h e 6 0 composition of the fluid may be checked f rom an additional analysis of coexisting qua r t z and use of t h e quartz-H 0 r c t ion (Table 6.3, eq. (3)). In addition, having t i e 6"O values of both qua r t z and se r i c i t e allows use of t h e quartz-muscovite isotopic geothermometer (Table 6.5, eq. (411, which thereby serves t o check t h e validity of t h e assumed temperature .


There a r e var iants t o the foregoing procedures, and where applicable they may offer advantages in t e r m s of t i m e and expense. Provided i t can be assumed t h a t t he hydrothermal fluid was predominantly of me teo r i c origin and t h a t t h e system was character ized by large water-to-rock ratios, i t is possible t o e s t ima te both t h e hydrogen- and oxy en- isotopic composition of the fluid from a single "0 analysis of qua r t z or o ther oxygen-bearing mineral. This approximation is feasible, t o the e x t e n t t h a t a l l pfgceding assumptions a r e valid, because the D and 6 0 contents of meteor ic waters, although extraordinarily variable, change sympathetically and linearly according t o t h e equation

c i t ed by Craig (1966) and Taylor (1974a). However, because of t he qualification with respect t water:rock ratios, which if low may per turb the 61'0 compositions of fluids and minerals precipitated t erefrom, i t is normally customary t o ca lcula te t h e 61'0 values of t he fluids f rom the 6 D values of minerals using equation (12), or from measured 6 D values of fluids ex t r ac t ed from mineral inclusions. Me o r i c waters a r e character is t ica l ly depleted in D and "0 by values t h a t may exceed -150 and -20 permil, respectively, r a t ive t o thei r oceanic source (0 O/oo for 6 D and Sf4 0, by definition). These depletions t a k e place a s a result of equilibrium fractionations a t low t empera tu res t h a t accompany the progressive condensation and crystall ization of rain and snow from a f in i te quantity of a tmospher ic water vapor (see Figs. 6.2 and 6.3, curves 1 and 12, respectively). Their magnitudes vary directly with alt i tude, la t i tude , and t h e re la t ive amount of water vapor removed from t h e system. Because isotopic studies have demonstra ted t h a t waters of meteor ic origin dominate geothermal and most hydrothermal systems, especially those of epi thermal character , t he present-day compositions of these waters over much of North America is shown in Figure 6.6 (af ter Taylor, 1979, p. 243). Note t h a t precipitation over Nevada, and most of the Basin ?dld Range province, is character ized by 6 D and 6 0 values of -130 t o -80 and -80 t o -11 permil, respectively. These depletions a r e similar t o those of "fossil" waters calculated from the mineral da t a for Ter t iary hydrothermal sys tems of this region. How- ever , i t is appropriate t o conclude this discussion of "calculated" fluid compositions with a no te of caution. Studies by Truesdell (1974) and o the r s and the d a t a and discussions of Friedman and O'Neil (1977) and Taylor (1967, 1979) indicate t h a t t h e ca lcula ted compositions of water from saline hydrothermal l u ~ d s may be in er ror , and usually depleted in D and "0 with respect t o t h e t r u e values. This solute e f f e c t is caused by t h e tendency of some cations, particularly those of l a rge ionic potential , t o hydrate in solution with t h e isotopically heavier molecules of water , which thus leaves the f r e e water t h a t equilibrat with minerals proportionately depleted in D and "0. Although difficult t o evaluate , this solute-controlled isotopic e f f e c t is probably small in epi thermal sys tems having fluids of low salinity.

Figure 6.6. Distributions of 6 D and 6180 in meteoric waters over part of North America (after Taylor, 1979) .

Other applications of t h e s table isotopes t o investigations of provenance warrant brief mention. It i s well known t h a t authigenic minerals formed in sedi- men ry environments a r e conspicuously enriched in '$0 rela t ive t o their magmat i c counterpar ts because of t he larger f rac t ionat ions permit ted a t low te eratures. Smaller, ye t significant, enr ichments of7$ found in some igneous rocks and minerals have been used in conjunction with o ther geochemical and geologic d a t a t o document examples and sources of rnagma contamination and t h e sedimentary component of S-type grani tes (Magaritz e t al., 1978, and references therein). The isotopes of carbon and sulfur also undergo appreciable fractionations a t low t o in termedia te temperatures , and these may be enhanced when t h e react ion pairs involve both oxidized and reduced forms of these e lements (Fig. 6.1, curves 1, 2, and 3; Fig. 6.4, curves 1 and 2). As a

nsequence, t h e carbonates a r e commonly enriched in I3C relative t o graphite, organic carbon, hydrocarbons, and marine sul fa tes a r e enriched in re la t ive t o magmat i c sulfur ( " 0 O/oo) and in contras t t o 34~-dep le t ed sedimentary sulfides. These enrichments and depletions originally served a s a qualitatively convenient means by which to in terpre t t he isotopic da ta in t e rms of oceanic, magmat ic , and biogenic sources and processes. However, Sakai (1968) and Ohmoto (1972) demonstra ted the fallibility of subjective in terpre ta t ions by showing t h a t t he isotopes of carbon and sulfur could undergo large fractionations a t relatively high t empera tu res by way of inorganic redox reactions controlled by Eh and pH. In spi te of

errors inherent with i sotopic generalizations, Thode et al. (1954) and Feely and Kulp (1957) deduced from carbon- and sulfur-isotopic d a t a the biogenic origin of ca l c i t e and nat ive sulfur in t h e cap rock of Gulf Coast s a l t domes f rom sources of evaporit ic anhydrite a t depth. In addition t o t h e sc ient i f ic mer i ts of thei r contribution, t h e pract ica l corollary was t h a t s a l t domes lacking a n anhydrite-calcite cap rock a r e unlikely t o conta in economic accumulations of nat ive sulfur and petroleum!

GEOLOGIC DISTRIBUTIONS

Abundances of t he s t ab le isotopes in geologically important environments a r e now summarized and briefly described. The purposes a r e twofold: f irst , these d a t a o f fe r background and perspective for discussions of t h e epi thermal environment t h a t follow; and second, t hey exhibit trends t h a t a r e largely consistent with those derived from theory and exper iment a s previously described. For some readers i t may be disconcerting t o observe t h a t t h e isotopic range fo r any of t h e four e l emen t s may be broad and overlapping f rom one environment t o another. However, th is appa ren t lack of isotopic definition is principally an a r t i f a c t of these compilations. The isotopic signature of a n e l emen t for a particular member of an environment, such a s a single mineral deposit , plutonic phase, sedimentary formation, etc., is usually narrow and well defined. Nonetheless, t he environment may exhibit a larger isotopic range because i t s members may b e compositionally dissimilar owing t o d i f ferences of provenance, depositional conditions, and age.

Hydrogen and Oxygen

I t is evident f rom previous discussions t h a t D and ''0 form an isotopic couplet t h a t is uniquely suited t o the study of fluids and minerals in aqueous systems. However, because oxygen is a major const i tuent of t h e c rus t ("46.6 wt%), whereas hydrogen (20.1 wt%) is not, t h e ''0 con ten t of fluids is more likely t o be per turbed by water-rock reactions than is t he D content. Accor ngly, i t is useful t o f i rs t review t h e distributions of "0 in various geologic environments. These da ta , a s por t rayed in Figure 6.7, a r e taken largely from the summaries of Taylor (1967, 1974a, 19791, Garlick (1972 Faure (l977), Hoefs (19801, and 1% other sources. The6 0 values for ultramafics ( > 5 and < 7 O/oo) a r e similar t o those of meteor i tes (-3 t o 7'/00) and a r e consistent with t h e presumed man t l e origin of t h e former . Mor siliceous igneous clans a r e progressively enriched in "0 for the sequence f rom basal ts and gabbros (-5.5 t o <8 O/oo), through andesi tes and granodiorites (-5.5 t o >I2 O/oo), t o rhyolites and grani tes ("6 t o 1 3 O/oo). This generalized isotopic t rend is compat ib le with observed f rac t ionat ions between t h e common rock-forming minerals wherein 6180 values a r e larges t in quartz, carbonates, and alkali feldspars; in termedia te in plagioclase feldspars, micas, and ferromagnesian minerals; and smalles in t h e Fe-Ti oxides. A similar trend of diminishing 5'0 values is evident among the authigenic minerals of

mar ine sedimentary rocks t h a t include che r t s (-20 t o 39 O/oo: Garlick, 1972; Kolodny and Epstein, 19761, carbonates (-15 t o 36 O/oo: Garlick 1972; Veizer and b Hoefs, 19761, shales (- 11 t o 29 loo: Savin and Epstein, 1970b), and ferromanganese dules ("10 t o I4 O/oo; Field et al., 1983). The largel'O enrichment of sedimentary rocks and minerals a s compared t o those of magmat i c origin results f rom the larger f rac t ionat ions permit ted a t t h e low t empera tu res prevailing in the hydrosphere, and in spi te of t he f a c t t h a t magmas a r e enriched in ''0 (-6 t o 10 O/OO) re la t ive t o s e a water ("0 O/oo). This distinction between isotopically heavy authigenic sedimentary minerals and thei r lighter magmat ic counterpar ts gives support t o hypotheses of assimilation or a n a t e c t i c mel t ing of sedimentary r ks t o account for t h e f ew documented examples of "0-enriched (up t o 16 O/oo) plutonic and volcanic rocks (Taylor and Turi, 1976; Magaritz e t al., 1978). By analogy, t h e sources of fluids and mineral consti tuents in hydrothermal sys tems may also be constrained by similar isotopic differences. The d a t a fo r hydrothermal deposits in Figure 6.7 a r e subdivided between Cordilleran and volcanogenic massive-sulfide types, with t h e principal d i f ferences being t h a t t h e former a r e associated with epizonal plutons whereas the l a t t e r a r e deposited in a submarine environment on or shor t distances below the s e a floor (Sawkins, 1972). For purposes of comparison, we have excluded t h e d a t a for epi thermal deposits f rom t h e Cordilleran subgroup a s su mar ized in

18" Figures 6.7, 6.9, and 6.10. Values of 6 0 for qua r t z (2-4 t o 1 3 O/oo) carbonates (<6 t o 14O/oo), and d sul fa tes (-6 t o 20 /oo) of t h e Cordilleran subgroup a r e chiefly from Sheppard e t al. (19711, Fuex and Baker (19731, and Watanabe and Sakai (1983). Although the re a r e abundant d a t a fo r hydrothermal s i l ica tes o ther than qua r t z (feldspars, micas, clays, oxides, etc.), they a r e not i l lus t ra ted in Figure 6.7 t o preserve c lar i ty among t h e dominant mineral phases. Their isotopic distributions would largely mim' those of quartz, but they would be more depleted in '$0 because of smaller f rac t ionat ion f ac to r s a t any given t e m p a t u r e (Table f5 6.3 and Fig. 6.3). Distributions of 6 0 values in qua r t z ( -7 t o 14 O/oo), carbonates (=9 t o 20°/oo), and sul fa tes ("5 t o 15 O/oo and more) of t h e volcanogenic massive-sulfide deposits a r e taken f rom Kusakabe and Chiba (19831, Watanabe and Sakai (1983), and Fifarek (1985). Although t h e d a t a base for massive-sulfide deposits is less extensive than t h a t fo r t h e Cordilleran types, the i r minerals appear be isotopically less varlable and more enriched in '0. Such differences, if real , probably r e l a t e t o the larger range of depositional t empera tu res and differing sources of t he aqueous component (magmat ic versus meteor ic) in Cordilleran hydrothermal systems.

D a t a re levant t o the foregoing iscussion a r e 1% given by t h e distributions of dD and 6 0 in various

waters and minerals, a s displayed in Figure 6.8, provide d a t a re levant t o the foregoing discussion, and a r e i l lustrative of several useful applications of t he hydrogen-oxygen isotope pair. The isotopic locations of s tandard mean ocean water (SMOW) and t h e me teo r i c wa te r line (MWL) a s derived from equation 12 (Craig, 1966; Taylor, 1974a), plotted a t t he top cen te r and diagonally down the left-hand margin,

Environment

nereorlres

1gneovs R o c k s

ulrramafr~s

basaits/gabbros

andeaiLes/granod~orires

rhyol,tes/graniten

Sedimentary Rocks

chert=

Carbonares

shales

sulfates

Fe-"n nodules

Hydrothemal Deposrtr

Cordilleran

quartz

carbonates

sulfates

Vols. Hass. Sulfide

quartz

carbonates

Sulfates

Figure 6.7. Distributions of 6180 in geologic environments.

respectively, a r e useful points of reference. The diagonal a t t h e right-hand margin represents t h e kaolinite l ine (KL) c i t ed by Sheppard e t al. (1969) and based on t h e work of Savin a n l f p s t e i n (1970a), which marks a continuum of 6 D and d 0 values in kaolinites f rom modern soils. Parallelism of lines MWL and KL documents equilibrium isotope exchange of D and 1 8 0 between me teo r i c waters and kaolinites formed during t h e weathering and transformation of rocks t o clay- rich soils. The compositional domain of magmat i c waters, given by the rectangular box a t center-right of Figure 6.7, has been calculated from isotopic analyses of hydrous magmat i c silicates. Most magmat i c waters have a relatively confined range of values e tween -85 and -40 permil D and 5.5 t o 9 permil 6 & 0 (Taylor, 1974a; 1979). Not i l lustrated, for reasons of c lar i ty and possible lack of relevance, a r e the broad isotopic fields of metamorphic waters and t h e saline brines of sedimentary basins. According t o Taylor (1979), t he metamorphic waters inherit the i r compositio a1

I 6 variability ('-65 t o -20 O/oo AD, and 5 t o 25 '100 6 0 ) f rom dehydration and fluid-mineral react ions with isotopically variable igneous and sedimentary rocks, whereas t h e d a t a fo r t he brines a r e broadly sca t t e red t o t h e right of t h e MWL and suggest varying mixtures of both connate and meteor ic waters t h a t have been modified by other fluid-sediment in teract ions a t depth within t h e basins.

Compositional variations of waters associa ted with a number of well-known geothermal systems, a s modified f rom the d a t a of Craig (19661, White e t al. (19731, and Truesdell and Hulston (19801, a r e p lot ted on Figure 6.8 Surface waters (large closed circles) have 6D and 6180 values located on or near the MWL, whereas t h e values for re la ted subsurface wa te r s (small closed circles) trend variably and horizontally t o t h e right (dashed lines) from the MWL. Such trends, known a s t h e "oxygen isotope shift", a r e cha rac te r i zed by increasing values of 6180 a t nearly constant 6 D and a r e common to the subsurface fluids of many

Figure 6.8. Distributions of 6 D and 6 180 in various waters, minerals, and hydrothermal fluids of epithermal deposits.

geothermal systems. The shi f t t o larger 6180 values i s caused b equilibrium isotope-exchange r e a c i ns between r80-depleted me teo r i c waters and "0- enriched rocks during I%ater-rock reactions. In general, t he s ize of t he 0-shi f t corre la tes d i rec t ly with t empera tu re and salinity of t h e fluids, and inversely with t h e mass ra t ios of wa te r t o rock. Because t h e e f f ec t s of t he 180-shift and water:rock ra t ios a r e in ter re la ted and may influence in terpre ta t ions of t h e analytical d a t a for fluids and minerals, these phenomena will be discussed a t g rea t e r length in forthcoming considerations of t h e epi thermal deposits. However, t he reader should no te t h a t me teo r i c waters similar t o those within t h e compositional range from Wairakei/Broadlands t o the Sal n Sea (6D 2-40 t o -80 O/oo) could be driven by a n '%-shift in to t h e isotopic domain of magmat i c wa te r s by exchange reactions involving smal l t o subequal amounts of water re la t ive t o t h a t of rock (low t o in termedia te water:rock mass ratios). Values of AD, in contras t t o those of 6180, a r e not significantly a f f ec t ed by water:rock exchange because a t t hese water:rock mass ra t ios t h e principal source of hydrogen is in t h e aqueous fluids. Thus, t h e geothermal fluids re ta in the 6D value of thei r me teo r i c source, and water-rock isotope-exchange react ions a r e manifest a s the subhorizontal t rend lines of Figure 6.8. The slight positive slope t o a f e w t rend lines, such a s fo r t he Salton Sea and Yellowstone geothermal areas, may result from deuter ium enr ichment t h a t i s unrelated to water-rock reactions. Concentra t ion of deuterium is caused by evaporation of surface waters prior t o recharge a t t h e Sal ton Sea (Craig, 1966) and by t h e boiling of subsurface wa te r s a t Yellowstone (Truesdell e t al., 1977) and perhaps elsewhere.

As previously noted, t h e isotopic compositions of hydrothermal and other f l i s may be calcula ted f rom 1 .f t h e measured 6 D and 6 0 values of associa ted minerals by means of t h e appropr ia te mineral-water

f rac t ionat ion curves (Tables 6.2 and 6.3) and using an e s t ima ted t empera tu re of deposition. The "calculated" fluid compositions derived f rom many bioti tes and a f ew ser ic i tes of porphyry-type deposits plot within the magmat i c water box (Sheppard e t al., 1971; Osatenko and Jones, 1976; Taylor, 1979). However, t he d a t a for a f ew biotite, most ser ic i tes , and nearly a l l o ther hydrous minerals of h y d r ~ t h e r m a l ~ g r i g i n a r e located within t h e broad range of 6 D and 6 0 values between t h e MWL and K L boundaries. Compositions of t h e fluids ca lcula ted from the isotopic d a t a for these minerals, or from t h e fluids of inclusions contained therein, suggest wa te r s of e i ther me teo r i c origin, or of mixed meteor ic-magmat ic or magmatic-oceanic parentage.

According t o Sheppard e t al. (1969) and Taylor (1974a), minerals such a s t h e kaolinites t h a t may form ei ther by hypogene o r supergene processes may be distinguished on t h e basis of temperature-controlled fractionations, which de te rmine thei r isotopic positions re la t ive t o t h e hypogene/supergene kaolinite line of Figure 6.8. Supergene kaolinites formed a t low temperatures and in equilibration with me teo r i c waters plot t o the right of this l ine and up t o the KL, whereas the hypogene kaolinites formed a t higher temperatures plot t o t h e lef t .

Also i l lustrated on Figure 6.8 a r e t h e l 'calculated" compositions of hydrothermal fluids responsible for t he deposition of many epi thermal precious-metal deposits (closed triangles) of the western U.S. and a f ew elsewhere. These d a t a a r e f rom the results of o ther investigators and will be c i t ed subsequently. They a r e based on t h e 6 D values obtained from ?Ifid inclusions and(or) hydrous minerals, and on 6 0 values determined from qua r t z and(or) other associa ted minerals. The broad isotopic distribution of these "calculated" compositions largely precludes a significant input of magmat i c water t o these hydrothermal systems, but i t does record a major contribution from meteo r i c sources. Fur ther detail and elaboration about these fluids will be deferred t o our concluding discussion of the epi thermal deposits.

Carbon

Abundances of I 3 c in various geologic environments a r e i l lustrated in Figure 6.9, and they a r e based largely on t h e d a t a c i t ed by Cra ig (19531, Bender (1972), Fuex and Baker (1973), Ohmoto nd Rye

14 (1979), and Hoefs (1980). Values of 6 C a r e surprisingly variable in the products of high- t e m ~ e r a t u r e environments such a s me teo r i t e s ( - - I2 t o 9 Ojoo), igneous rocks ('-10 t o 3 O/oo), diamonds ( -6 t o 3 O/oo), and carbonat i tes (2-10 t o 2 O/oo); particularly because t h e compositional ex t r emes have been omi t t ed from these ranges. The causes of such isotopic variability a r e uncertain, but possibly r e l a t e to equilibriurn or kinetic redox reactions, contamination, inhomogeneities of source, and(or) differing proportions of compositionally distinct carbon in t h e samples. For example, t h e d a t a given for meteor i tes and igneous rocks ' s t h a t of t o t a l carbon, which consists both of "C-enriched and oxidized (carbonate, up t o 66 '100) and 13c-depleted and reduced (graphite, "organic," carbonyl, etc.; up t o

-30 O/oo) forms of carbon. In accordance with f rac t ionat ion theory (Table 6.1 and F' 6.1), similar re la t ive enr ichments and depletions of "C a r e present between t h e oxidized (CO -10 t o 2 '100) and reduced (CHb = -31 t o -16 ~ /o$ ' componen t s of volcanic/ geot ermal gases. Narrower compositional variations cha rac te r i ze mar ine l imestones (2-5 t o 4 O/oo), s e a water H C O (-5 t o -2 O/oo), a a tmospher ic C 0 2 ("-8 t o -6 O/oo), and the re la t ive "C enrichments among these compounds (CaCo3> HCO?COZ) a r e consistent with experimental-theoretical f r a c t ~ o n a t i o n trends. The c lus ter of values around and near 0 permil is expectable because t h e carbon PDB s tandard is ca lc i te of mar ine derivation. Fract ionat ions t h a t accompany kinetic photosynthetic react ions of a tmospher ic and hydrospheric C 0 2 t o form organic carbon lead t o marked depletions of 1 3 c in mar ine and land plants ('-34 t o -12 O/oo), and th i s isotopic record of biogenic processes i s preserved in carbonaceous sediments, coal, and petroleum ('-35 t o -10 '100). Although the pronounced isotopic d i f ferences between organic carbon and inorganic carbonates seem academic in view of obvious visible d i f ferences between these compounds, they do se rve a s a t r ace r of biogenic precursors where reduced forms of carbon become oxidized and remobilized in some magmatic, metamorphic, and hydrothermal (?) environments. The da ta for hydrothermal carbonates provided by Sheppard e t al. (19711, Fuex and Baker (19731, Ohmoto and Rye (19791, and Fifarek (1985) show remarkably l i t t le isotopic variabili ty (--I 1 t o 1 O/oo), regardless of textura l var ie ty or gene t i c occurrence. According to Ohmoto and Rye (19791, the 613c value of carbon in mantle-derived igneous rocks is -5 * 2 permil, and this value is not demonstrably di f ferent from t h a t of carbon in average sedimentary or crus ta l rocks based on considerations of mass balance. Thus, carbon isotopes do not offer promise a s a means for distinguishing between mant le and crus ta l sources of magma and igneous rock. The carbon in minerals and fluids of hydrothermal sys tems may be derived from mant le sources, a s i s permissive from the isotopic evidence, o r from diverse sources in country rocks e i ther by oxidation of reduced forms o r by dissolution and(or) decarbonation reactions of carbonates (Ohmoto and Rye, 1979).

Sulfur

Isotopic abundances of sulfur portrayed in Figure 6.10 a r e taken principally from t h e d a t a and summaries of Field (19721, Field e t al. (1976, 1983, 19841, Ohmoto and Rye (1979), Claypool e t al. (1980), Hoefs (19801, and Sakai e t al. (1984). Compositions of or thomagmat ic to t a l sulfur in igneous rocks (2-3 t o 3 O/oo) a r e predictably close t o t h a t of meteor i t ic sulfide-sulfur (20 O/oo). However, those of t he component oxidized and 34~-enr i ched sul fa te (up t o 10 '100) and reduced and 3 4 ~ - d e p l e t e d sulfide (up t o -10 '100) fo rms a r e more variable, because of redox reactions. The 6 3 4 ~ values of magmat i c Cu-Fe-Ni sulfides in layered maf i c intrusions (--6 t o 14 '100) exhibit larger variations a t t r ibutable both t o contamination from sedimentary sources in nearby country rocks and t o redox reactions within the host

Environment

ueteorites: total-c

Igneous ROC*.: total-C

carbonatites

diamonds

Volcanic/Geothemal

carbonate*

=O2

CH4 Graphite

A-spheric C02

sea water "CO;

~ilneetones lmrine)

organic Carbon

marine plants

coal, petrolem, and ~arbon.CeOYB matter

n i s ~ i s s i p p ~ valley: caco3

Hydrothermal: C.C03

PorPhYry-tyPe

vein-type

replacement-type

"01C. ma*.. sulfide

Figure 69. Distributions of 613c in geologic environments.

magma chambers. Sulfur-bearing products of volcanic/geothermal emanat ions a r e broadly isotopically similar to, but more variable, than magmat i c sulfur (=-3 t 3 O/oo), and they a r e increasingly deple ted in 3'S with some overlap in t h e redox sequence f rom SO (=-8 t o 18 O/oo), through native sulfur (2-15 t o 3 6 O/oo), t o H S (--9 t o 6 O/oo). Sea wa te r su l f a t e has a value of '20 *-I O/oo a t present, but has ranged f rom about 10 to 30 O/oo over Phanerozoic t i m e a s deduced f rom studies of m rine evaporites (Claypool et al., 1980). Thus, t he 3eS-age curve serves within broad l imits t o d a t e

mar ine sedimentary s t r a t (and volcanogenic massive- sulfide deposits). The "S-enriched sul fa te (=20 t o 35 O/oo) and sulfide ('-6 t o 25 O/oo) minerals of t h e and(or) evaporites. Sedimentary sulfides of diagenetic-syngenetic or l a t e r epi e n e t i c origin have 5 an extraordinarily large range of 6 4~ values (2-50 t o 50 O/oo). Such large variations, especially t h e 3 4 ~ depletions, have been documented by exper imenta l investigations, and they a r e caused by kinetic fractionations of -20 t o -50 permil and more t h a t may accompany t h e biogenic reduction of SO; t o H S (Ohmoto a n g p y e , 1979). Although sulfides variably depleted in S a r e considered to be typical of t h e formed by biogenic processes, o thers may have "S enr ichments t h a t r e l a t e t o f ac to r s such a s source, reservoir of sul fa te , and other conditions within the system. Regardless of environment or locale, hydr thermal sul fa tes of hypogene origin a r e enriched in $'S re la t ive t o associa ted sulfides, which is consistent with f rac t ionat ion theory and the inferred range of depositional t empera tu res f rom about 2 0 0 ' ~ t o 6 0 0 ' ~ (Figs. 6.4 and 6.5). Contrary t o t h e impressi n given by t h e d a t a in Figure 6.10, variations in the a 4 S values fo r e i ther sul fa tes o r sulfides of individual deposits rarely exceed 5 to 7 O/oo, although some deposits may exhibit a compositionally distinct cluster of absolute values t h a t may re l a t e t o source, age, and(or) unique conditions within t h e system (Field e t al., 1983; Fifarek, 1985). Both t h e sul fa tes and sulfides of Cordilleran-type hydrothermal deposits a r e

generally deple ted in 3 4 ~ re la t ive t o thei r counterpar ts in volcanogenic massive-sulfide deposits. The Cordilleran subgroup includes a large spectrum of porphyry, vein, and replacement types of deposits t h a t apparently a r e devoid of a distinctive isotopic signature regardless of d i f ferences in host rock, metals, minera and depositional textures of o r e and gangue. The '% d a t a for sul fa tes and sulfides of a few deposits have values suggestive of country rock contamination, but t h e majority a r e consistent with derivation f rom a -0 permil source of deep-seated "magmatic" sulfur. In contras t , su l fa tes and sulfides of t h e v lcanogenic massive-sulfide subgroup a r e normally "S-enriched because they derive thei r sulfur largely o r ent i re ly f rom isotopically heavy sea-water sulfate. Da ta fo r t h e sul fa tes (= 12 t o 39 O/oo) a r e essentially equivalent t o those of temporally similar mar ine evapor i t ic sul fa tes (-10 t o 30 O/oo), and those for associated sulfides (28 t o 22 O/oo) a r e corre la t ive t o the age-trend but deple ted in 3 4 ~ by about 15 t o 18 O/oo re la t ive t o oceanic sulfate-sulfur a s a consequence of t empera tu re dependent fractionation (Sangster, 1968; Franklin e t al., 1981; Fifarek, 1985). The da ta for a single mid-ocean hydrothermal vent (21' North, Eas t Paci f ic Rise, Baja, California) a r e i l lustrative of such isotopic e f f ec t s under contemporary oceanic conditions (Styrt e t al., 1981).

EPITHERMAL DEPOSITS

Distributions of t h e s table isotopes in epi thermal deposits a r e now considered. Also included a r e the d a t a for severa l geothermal sys tems because they a r e regarded by many t o be contemporary analogues of t h e epi thermal environment (White, 1981). For t h e purposes of comparison and discussion, we have subdivided the epi thermal deposits in to sediment- hosted, volcanic-hosted, and zoned polymetall ic vein occurrences. Our subdivisions d i f fer partly from those of Hayba e t al. (1985, this volume) in t h a t most o r a l l deposits of our zoned polymetall ic vein subtype a re , or may be, grouped in thei r Adularia-Sericite subtype of

0 5 /w

environmenr -10 -10 O +lo 2 0 t10 +,P

4 Meteorire. ---- rgneous R O C ~ ~ : r o t a ~ - $ -------

Cu-Fe-Ri aulfides of layered mafic 1ntr.

"olsanis Em.n.ci.ns

native sulfur

=zS sea Wakar 30; - Sedimentary Rocks

sulfates - sulfides - 5 0 ---------.+so

missi..lppi valley Deposits

sulfates

sulfides

HydIOrheFD.1 Deposit.

Cordillar.": ."Ifate*

avlfides

"010. msa: sulfates

sulfides

Ocean I~dge: sulfatea - sulfldell -

Figure 6.10. Distributions of 6 3 4 ~ in geologic environments.

volcanic-hosted epi thermal deposits. These deposits may exhibit character is t ics suggestive of a deeper level of hydrothermal mineralization, such a s fluid inclusions having higher salinit ies and homogenization temperatures ; a not uncommon magmat ic component t o t h e hydrothermal fluids a s deduced f rom hydrogen- and oxygen-isotope data ; host rocks t h a t include plutonic and sedimentary lithologies; a l tera t ion t h a t lacks widespread and pervasive zones of advanced argil l ic and alunitic assemblages; and o res t h a t contain relatively abundant sulfides and sulfosalts of t h e base metals. Presumably such differences a r e not genetic, but ins tead represent variations in style and content of mineralization t h a t have developed from chemical and the rma l gradients in deeply convecting hydrothermal systems.

Carbon

D a t a for 1 3 c in various geothermal sys tems and e ' t he rma l deposits a r e i l lustrated in Figure 6.1 1. The (PJC values for a l l occurrences, excep t t h e Geysers and Pueblo Viejo, a r e remarkably uniform within t h e narrow range of -10 t o 1 permil, and suggest t h a t t h e carbon was derived e i ther from magmafjic or mar ine l imestone sources (Fig. 6.9). Extreme C depletions (-25 t o -24 O/oo) of carbonaceous mater ia l in volcaniclastic sedimentary rocks a t Pueblo Viejo (Kesler e t al., 1981) a r e typical of reduced fo rms of organic carbon. Oxidized carbon cornpounds such a s C 0 2 , aqueous HCO, host-rock carbonate, and vein ca l c i t e a t t h e Geysers geothermal a r e a exhibit la rge variations in613c (-19 t o 1 O/oo) according t o t h e work of White e t al. (1973) and Sternfeld (19811, and other references c i ted by these authors. Detailed fluid- inclusion, isotopic, and mineralogical Sternfe ld (1981) suggest t h a t much of the variabili ty may be a t t r ibu ted t o multiple sources of biogenic, marine, and magmat i c carbon in t h e igneous and sedimentary host rocks t h a t were rernobilized and redeposited during subsequent and temporally sepa ra t e metamorphic and geothermal events. Vein calc i tes diminish in 613c ( ~ 4 O/oo/lOOO m) with increasing depth, a s a result of temperature-ind fractionation, and a distinct population of deple ted calc i tes is a t t r ibu ted t o late-stage re- equilibration with C02-r ich s team. In contras t t o t h e Geysers, isotopic compositions of carbon in carbonate c las ts of host rocks a t t h e Cer ro Pr ie to (Williams and Elders, 1984) and Salton Sea (Clayton et al., 1968) geothermal a reas and in vein carbonates f rom these and t h e Broadlands (Blattner, 1975) and Wairakei (Clayton and Steiner, 1975) reas of New Zealand a r e

I 4 less variable. Values of 6 C in host rock and vein carbonates decrease with increasing depth a t Broadlands and Cer ro Prieto. The cause of these isotopic t rends is uncertain. They may possibly r e l a t e t o diminishing equilibrium fract ionat ions with increasing temperatures a t depth, and (or) t o o ther complexities such a s boiling, d carbonation and (or) dissolution reactions, influx of 'JC-depleted organic carbon, and possibly other processes (see Blattner, 1975; Williarns and Elders, 1984). Da ta for carbonates from Wairakei (Clayton and Steiner, 1975) and t h e Salton Sea (Clayton e t al., 1968) do not show isotopic

trends re1 t e d t o depth o r reservoir temperatures . However, I3C-depleted carbonates f rom t h e Salton Sea field corre la te inversely with t h e t o t a l carbonate (wt.%) content of t h e host, which suggests they a re , the residuals of decarbonation react ions accompanied by t h e loss of 13c-enriched C 0 2 (Clayton e t al., 1968).

Carbonates from unaltered and a l t e red host rocks and veins of t h e sediment-hosted epi thermal deposits a t Carlin (Radtke e t al., 1980) and C o r t e z (Rye e t a1 1974) exhibit a narrow and overlapping

1'3 range of 6 C values (Fig. 6.11; %-6 t o 1 O/oo). These and other d a t a suggest t h a t most of t h e carbon in hydrothermal ca lc i tes was probably e x t r a c t e d by dissolution reactions from t h e carbonate host rocks a t depth. However, t he fluids a t Carlin a1 must have contained a component of oxidized '%-depleted organic carbon from the host rocks, provided isotopic equilibrium prevailed, t o account for t h e relatively light compositions of one main-stage and severa l l a t e low-temperature vein ca lc i tes (Radtke e t al., 1980).

The various carbonate minerals formed in zo polyrnetallic veins also show a narrow range of 6PPCd values ("J-10 t o 0.1 O/oo), but a s a group they a r e slightly deple ted in 1 3 c re la t ive t o those of sediment- hosted deposits. D a t a fo r rhodochrosites, manganosiderites, and s ider i tes f rom Creede, Colorado (-8.2 t o -4.0 O/oo; Bethke and Rye, 19791, ca lc i tes and other carbonates (?) f rom Tui, New Zealand (-7.8 t o O.lO/oo; Robinson, 19741, and ca l c i t e s from Casapalca, Peru (-10.0 t o -2.6 O/oo; R y e and Sawkins, 19741, a r e collectively in terpre ted a s being indicative of magmat i c carbon (%-5 2 2 O/oo; Ohmoto and Rye, 1979). This conclusion is also supported by hydrogen and oxygen-isotope d a t a of inclusion fluids and host minerals a t Creede and Casapalca. However, t he source of carbon in ca lc i tes and rhodocrosites a t t he Sunnyside mine (-7.9 t o -1.8 O/oo), near Creede, is not definit ive, and may have been derived in p a r t f rom dissolution of mar ine l imestones o r magmat i c sources a t depth, or from atmospher ic C O dissolved in circulating me teo r i c water ( ~ a s a d e v a l ? and Ohmoto, 1977).

Geothermal Systems

Geysers: ::;- who?e rock calcite

salton Sea/Cerro Prieto: carbonate host altered carb. host

carbonates

Broadlands/wairakei: calcite

Sediment-Hosted

Cortez: carbonate host altered carb. host

calcite

carlin: carbonate host altered carl). host

calcite

Volcanic-Hosted

Pueblo Viejo: carbonaceous sed.

Zoned Polymetallic Veins

Creede: carbonates

Sunnyside: carbonates

Tui: carbonates

Casapalca: calcite

Figure 6.11. Distributions of 613c in epither- ma1 deposits.


Thus, in the absence of contributing information from other isotopic e lements , fluid inclusions, nd assemblages of vein and a l tera t ion minerals, t he 'C d a t a alone a r e unlikely t o provide unique in t e rp re t tions t o geochemical problems. Moreover, ver t ica l I 3 C gradients present in some geothermal sys tems have not been repor ted in the epi thermal deposits. However, ea r ly calcites (-10.1 t o -6.1 O/oo) a r e deple ted in 1 3 c re la t ive to l a t e ca lc i tes (-6.4 t o -2.6 O/oo) a t Casapalca (Rye and Sawkins, 19741, and rhodochrosites (-7.9 t o -6.3 O/oo) a r e depleted in 1 3 c re la t ive t o ca lc i tes (-3.8 t o -2.8 O/oo) a t t he Sunnyside mine (Casadevall and Ohmoto, 1977).

Sulfur

Abundances of 3 4 ~ in geothermal sys tems and epi thermal deposits a r e portrayed in Figure 6.12. As expected, sulfide-sulfur from whole-rock samples of post-glacial basalts near the geothermal fields in Iceland vield 6 3 4 ~ values (-1.8 t o 0.4 O/oo) similar t o those O ~ O permil magmat ic sulfur ( ~ a k a i e t al., 1980). The d a t a for aqueous SO: and sul fa te minerals occupy t w o distinct isotopic populations. One group represented by samples f rom Iceland (Sakai e t al., 19801, Yellowstone (Schoen and Rye, 1970; Truesdell et al., 1977, 1978), and Wairakei (Steiner and Raf t e r , 1966) is enriched in 3 4 ~ ('15 t o 2 3 O/oo) and consists of hypogene sulfates t h a t isotopically equilibrated with H2S a t depth and a t relatively high t empera tu res (?300°C). The other group, also from Iceland, Yellowstone, and Wairakei, is isotopically m o r e variable and relatively depleted in 3 4 ~ (-6 t o 12 O/oo). t consists of mixed proportions of t h e deep

S4S-enriched hypogene sul fa te and shallow 3 4 ~ - depleted "supergene sul fa te formed by near-surf ce , non-equilibrium, and quant i ta t ive oxidation of 'S- deple ted H2S (Truesdell e t al., 1977, 1978). Isotopically l ~ g h t aqueous and mineral sul fa tes a r e not only common t o the acid-sulfate springs of geothermal areas , but they a r e also typical of sorne alunites and bar i tes of volcanic-hosted epi thermal deposits; particularly those associated with advanced-argillic a l tera t ion ( the acid-sulfate type of Hayba e t al., 1985, this volume). Native sulfur a t Yellowstone has formed chiefly by near-surface inorganic oxidation of H2S, a s is consistent with the com ositional similarities B between e lementa l (-8.4 t o 3.2 /oo) and reduced (-5.0 to 4.0 O/oo) sulfur compounds (Schoen and R e, 1970). 4 However, t h e relatively broad spread of d S values (8.4 t o 4.0 O/oo for so and H2S @ 0 O/oo) under e s sl ight fractionation upon separation in to d iscre te "S- depleted vapor and 34~-enr iched aqueous phases (Truesdell e t al., 1978). The sulfide minerals normally exhibit narrow compositional ranges t h a t a r e similar t o those of associated H2S, although the re may be some variability a t t r ibutable t o differing sources of sulfur both within and between the geothermal fields. For example, the pyrites from Iceland consist of t w o distinct isotopic populations: one is enriched in 3 4 ~ (2.2.9 t o 7.9 O/oo) and has been derived through reduction of isotopically heavy seawa te r sul fa te , and t h e other is depleted in 3 4 ~ (-4.6 t o 0.9 O/oo) and has originated from a magmat ic source (Sakai e t al., 1980).

Geothermal Systems

Iceland: magmatrc sulfur SO4 sulfates H2S pyrite

Yellowstone: SO4 sulfur H2S

Salton Sea: s u l f ~ d e s

Broadlands: sulfides

Wairakei: SO4 sulfates Hz5 sulfides

Sediment-Hosted -- Cortez: barite

diagenetic py sulfides

Carlin: barite diagenetic py sulfldes

Volcanic-Hosted

Tolfa: sulfates sulfides

Pueblo Vlelo: sulfates sulfur sulfides

Goldfleld: alunrte pyrite

Zoned Polymetallic Veins - Sari Juan Mountains:

Creede: barite 45

sulfides -

Sunnyside: sulfates - sulfides

RICO: sulfides

Ouray: sulfldes

T U I : barlte - sulfides

Guanaluato: sulfides ~n country rock

s u l f ~ d e s in volc. sulfides in ore ---------

Casapalca: sulfides

Flnlandia: barite sulfides

Western Cascades: sulfides

Golden Sunlight: barlte s u l f ~ d e s

Figure 6.12. Distributions of 6 3 4 ~ in epither- ma1 deposits.

Ranges of 6 3 4 ~ a r e narrow for pyr i te and Ag-Cu sulfides (-1.4 t o 3.0 O/oo) of t h e Salton Sea; pyrite, pyrrhotite, galena, and sphaler i te (1.4 t o 5.1 O/oo) of Broadlands; and pyr i te and pyrrhot i te (2.7 t o 6.8 O/oo) of Wairakei. The sulfide compositions a t the Salton Sea a r e a a r e compatible with, but not proof of, a magmat ic source of sulfur (White, 1968, 1974), whereas those a t t h e Dro dlands and Wairakei a r e slightly more enriched in 94S and suggest a crus ta l provenance e i ther by leaching and par t ia l reduction of sul fa te f rom basement rocks (Drowne et al., 1975) or by magma generation in the upper c r u s t (Steiner and Raf t e r , 1966). Although t h e d a t a base is meager , and excep t fo r two sphalerite-galena pairs from Broadlands, t he re is l i t t le evidence of complete isotopic equilibrium between sulfate-sulfide and sulfide-sulfide assemblages in the geothermal environment.

All sulfide and most sul fa te minerals in and near t h e sediment-hosted gold deposits a t Carlin (Radtke e t a1 1980) and C o r t e z (Rye e t al., 1974) a r e enriched in 34\ re la t ive t o their counterpar ts in geothermal sys tems and other epi thermal deposits. The evidence, particularly f rom Carlin, suggests t h a t most of t h e

116 CHAPTER 6

sulfide-sulfur in pyrite-galena-realgar-sphalerite- s t ibni te o r e (4.2 t o 16.1 O/oo) was derived f rom t h e hydrothermal remobilization of diagenetic pyr i te (1 1.7 t o 14.3 O/oo a t Carlin; 5.1 t o 11.4 O/oo a t Cor t ez ) in t h e sedimentary host rocks. However, t he source of sulfate-sulfur in bar i tes (27.8 t o 31.7 O/oo a t Carlin) is uncertain, and may have originated e i ther by hydrothermal solution of sedimentary bar i te in t h e country rocks, or by f rac t ionat ion tha t may have accompanied par t ia l oxidation of sedimentary sulfide- sulfur incorporated in t h e fluids (Radtke e t al., 1980). Regardless of origin of t h e sulfate-sulfur, i t is remarkable t h a t isotopic t empera tu res ca lcula ted from barite-pyrite values ( 2 5 0 ' ~ t o 3 0 5 ' ~ ) a r e reasonably consistent with those derived from the homogenization of fluid inclusions ( 1 8 0 ' ~ t o 365'~).

The volcanic-hosted deposits of Tolfa (Italy), Pueblo Viejo and Goldfield have geologic f ea tu res common t o many geothermal sys tems in t h a t thei r volcanic host rocks have been pervasively a l t e red t o alunite-kaolinite + pyrophyllite assemblages of advanced argil l ic a l tera t ion. According t o Field and Lombardi (1972) and Cor t ecc i e t al. (19811, isotopic similarit ies between a luni te and bar i te (1.9 t o 9.6 O/oo) and hypogene pyrite, cinnabar, and galena (3.4 t o 10.3 O/oo) a t Tolfa suggest t h a t these sulfates, and possibly marcas i te (-1.5 and -0.6 O/oo), a r e of supergene origin. However, t he sulfate-sulfide assemblages of Tolfa may consist of two isotopically distinct populations because the 34~-dep le t ed marcas i tes a r e associa ted with t h e l ightest alunites (1.9 and 2.5 O/oo). Perhaps thes minerals formed in surficial acid-sulfate pools f rom g4S-depplted H2S t h a t had separa ted with boiling of reservoir fluids a t depth, a s proposed by Truesdell et al. (1978) for some hot springs of t h e Yellowstone area. In contras t , most sulfates a t Pueblo Viejo a r e considered t o be of hypogene origin by Kesler e t al. (198 because t h e bar i te and a luni te a r e enr iched in "ki re la t ive t o pyr i te and sphaler i te (18.8 t o 21.6 O/oo versus -10.1 t o -3.5 O/oo). However, Jensen e t al. (1971) have documented alunites of both hypo ene (11.6 to 8 23.3 O/oo) and supergene (-2.5 t o 1.7 /oo) origin a t Goldfield; t h e l a t t e r group bein8 isotopically similar t o hypogene pyr i te (-2.8 to 2.4 loo) from which they derived thei r sulfur. Thus, isotopic and geologic evidence suggest t h a t t h e sul fa tes (aqueous and mineral) of many geothermal and volcanic-hosted environments a r e o supergene origin, and acquired thei r distinc ive '4.S-depleted sulfate-sulfur (as contras ted t o 14S-enri hed hypogene sulfates) by near- surface oxidation of 54S-depleted hypogene sulfide- sulfur (H2S and mineral; s e e Field, 1966; Schoen and Rye, 1970; Jensen et al., 1971; Field and Lombardi, 1972). The source of hydrothermal sulfur a t Goldfield is considered t o be magmat i c (Jensen e t al., 1971). In contrast , t h e source a t Pueblo Viejo is thought t o be a mixture of prist ine and biogenically reduced sul fa te (15 O/oo) f rom Cre taceous s e a water (Kesler e t al., 19811, and t h a t a t Tolfa may have been derived f rom gypsiferous Miocene-Pliocene mudstones t h a t underlie this young Pliocene-Pleistocene volcanic complex (Field and Lombardi, 1972). This l a t t e r in terpre ta t ion is supported by the work of Taylor nd Turi (1976) who

1 6 a t t r ibu te the exceedingly high 6 0 values (15.3 t o

16.4 O/oo) of Tolfa quar tz l a t i t e s and rhyolites t o magmat i c assimilation of 180-enriched argillaceous sedimentary rocks a t depth. Although f ract ionat ion e f f e c t s between hypogene sulfide and sulfate-sulfide minera ls of t he volcanic-hosted group appear t o be consistent with t h e a t t a i n y g n t of a t leas t partial isotopic equlibrium (with 6 S values of sul fa tes > sulfides and those of sulfides mostly in t h e order py >sl > gn), a more precise evaluation of these apparent t rends is difficult because the d a t a for contemporaneous o r spatially associated mineral pairs and t r ip le ts a r e not available.

The isotopic d a t a base for zoned polymetall ic vein deposits (Fig. 6.12) is large and detailed, particularly fo r those of t he Creede district (Bethke e t al., 1973; Bethke and Rye, 1979; Foley et al., 1982; Hayba et al., 1985, this volume) and the Sunnyside mine of t h e San Juan Mountains, Colorado (Casadevall and Ohmoto, 1977); Tui, New Zealand (Robinson, 1974); Casapalca , Peru (Rye and Sawkins, 1974); Finlandia vein, Colqui district , Pe ru (Kamilli and Ohmoto, 1977); mining dis t r ic ts of t h e Western Cascades, Oregon (Taylor, 1971, 1974b; Power, 1985; Field and Power, 1985); and Golden Sunlight mine, Montana (Por ter and R,jfley, 1985). Sulfides in most of these deposits have 6 S values e a r 0 permil t h a t contras t markedly with associa ted 3'S-enriched hypogene sulfates, such a s a t Creede (-4.1 t o 1.7 O/oo versus 19.8 t o 45 O/oo), Sunnyside (-6.3 t o 2.7 O/oo versus 15.3 to 22.9 O/oo), Tui (-2.4 t o 4.9 O/oo versus 16.0 t o 19.5 O/oo) and d Finlandia (-4.0 t o 1.5 O/oo versus 14.0 t o 14.1 loo). Moreover, compositions of sulfides t h a t a r e unassociated with sulfates, such a s those given by Jensen et al. (1960) for Rico (-0.6 t o 4.0 O/oo) and Ouray (-2.0 t o 1.9 O/oo) in t h e San Juan Mountains of Colorado, and by Rye and Sawkins (1974) for Casapalca, a r e also closely grouped near 0 permil. Because of t h e overall isotopic similarity of most sulfides in these deposits t o t h e 0 permil value of magmat i c sulfur, and other geologic and geochemical considerations, a magmat i c source of sulfur is advocated by t h e authors of investigations a t Rico, Ouray, Casapalca , and t h e Western Cascades, and possibly a t C r e e d e and Finlandia. However, the commonly assumed "genetic" equivalence of 0 permil sulfides and magmat i c (0 '100) sources of sulfur is a hazardous generalization when applied to hydrothermal environments. I t i s valid for sulfides only to the e x t e n t i t may b e assumed t h a t concentrations of reduced sulfur (H2S) a r e approximately equal t o those of t o t a l sulfur in these systems. However, under conditions of high fO2 and (or) low pH, a s may be inferred from t h e presence of oxide and sul fa te minerals in vein and a l tera t ion assemblages, par t of t h e H2S becomes oxidized t o SO;, and to t a l sulfur must then consist of both o x i d ~ z e d and redu components. Because of t h e large fractionation of between sul fa tes and sulfides and of mass balance considerations between oxidized and reduced forms of aqueous sulfur in t h e fluids, sulfide minerals b come

f 4 increasingly deple ted in 3 4 ~ re la t ive to the 6 S of to t a l sulfur in the system, a s ratios of SO5:H S increase with increasing f 2 and (or) decreasing Redox changes a s desc r ibe2 such a s increasing s t a t e s of oxidation with evolution of hydrothermal systems,

have been proposed by Robinson (1974) t o account for t he progressive 3 4 ~ depletion observed in the paragenet ic order of sulfide depos' ion a t Tui. This so- called Eh-pH control of mineral "S compositions was f i rs t defined by Sakai (1968) and Ohmoto (1972), and has been subsequently refined by Rye and Ohmoto (1974) and Ohmoto and R y e (1979). I t is for reasons of this Eh-pH control, a s deduced from mineralogical and geochemical evidence, t h a t isotopically heavy marine sulfates have been proposed a s sources of sulfur in the hydrothermal deposits a t Sunnyside (upper Paleozoic evaporites of % I 2 O/oo; Casadevall and Ohmoto, 1977) and Tui (Jurass ic sea wa te r of ~ 1 6 '100; Robinson, 19741, and in spi te of t h e f a c t t h a t compositions of associated sulfides (-6.3 t o 2.7 O/oo and -2.4 t o 4.9 O/oo, respectively) bracket t h a t of 0 O/oo magmat ic sulfur. Moreover, this control is implicit t o t h e in terpre ta t ion of a magmat i c source of sulfur a t t h e Golden Sunlight deposit (Porter and Ripley, 19851, although both sul fa tes (1.2 t o 5.9 O/oo) and sulfi e s (-15.8 t o -4.0 O/oo) a r e appreciably depleted in '4S relative t o t h e major i ty of d a t a for hypogene equivalents (Fig. 6.12; also s e e Field and Gustafson, 1976, Fig. 3, for a graphical portrayal of this control). In contras t , Gross (1975) has proposed a crustal source fo r sulfur and meta ls contained in vein sulfides (-19.5 t o -3.4 O/oo) of pyrite-sphalerite- galena-argentite o re f rom Guanajuato, Mexico, which a r e considered t o have been derived from metal-rich and sulfide-bearing (-16.6 to 6.3 O/oo) Mesozoic sedimentary country rocks by heated ground waters during Oligocene volcanic activity. Fractionation trends fo r sulfate-sulfide and sulfide-sulfide mineral pairs from t h e zoned polymetall ic veins a r e largely consistent with those predicted from theory and experiment (Tables 6.4 and 6.5, and Figs. 6.4 and 6.5). However, with the exception of a few sphalerite- galena pairs, from Sunnyside, Tui, Casapalca, Finlandia, and the Western Cascades, t he 3 4 ~ t empera tu re e s t ima tes a r e rarely consistent with those obtained by fluid-inclusion homogenization methods. This disparity implies t h a t isotopic equilibrium largely did not prevail in these sys tems for reasons t h a t may include low depositional temperatures , relatively rapid accen t of fluids and deposition of minerals, and possibly abrupt changes in SOG:H2S ratios with the ascent of fluids in to more oxidizing environments. The lack of equilibrium between sul fa tes and sulfides i s clearly expectable based on t h e discussions and da ta presented by Ohmoto and Rye (1979) and Ohrnoto and Lasaga (1982). Moroever, biogenic processes (Hayba e t al., 1985, this volume), a s previously su gested by Kesler e t al. (1981) t o account for t he J'S-depleted compositions of bedded sulfides a t Pueblo Viejo, in addition t o t h e e f f e c t s of chemical and (or) isotopic disequilibrium (Bethke et al., 1973) a r e not considered t o be responsible for t h e large 3 4 ~ enr ichments (19 t o 45 'loo) of sul fa tes a t Creede, Colorado.

Hydrogen and Oxygen

Distributions of ''0 in the host rocks and minerals of many geothermal and epi thermal occurrences previously described a r e portrayed in

Figure 6.13. Because these d a t a a r e voluminous and based on numerous investigations of variable detail , our discussion will focus principally on t h e major isotopic t rends and reasons thereof. More deta i led information may be obtained a s needed f rom t h e references cited. Although these d a t a exhibit a large

Geothermal Systems

Geysers: steam whole rock quartz calcite

salton Sea/Cerro Prieto: carbonate host altered carb. host

quartz carbonates

Broadlanddwairakei: volcanic host altered volc. host

quartz sinter calcite adularia

Sediment-Hosted

Cortez: carbonate host altered carb. host

quartz calcite hematite

Carlin: carbonate host altered carb. host sedimentary chert

jaeperoid iuaitz barite calcite

Others: adularia

Vol~anic-Hosted

Tolfa: volcanic host altered volc. host

chalcedony

Comstock: altered volcanics quartz

Goldfield: altered volcanics quartz

Tonopah: volcanic host altered volcanic host

quartz calcite K-feldspar

Bodie: volcanic host altered volcanic host

quartz K-feldspar calcite

Others: quartz adularia

Zoned Polymetallic Veins

San Juan Mountains: altered plutonics altered volcanics

Creede: quartz carbonates illite chlorite

Sunnyside: quartz carbonates

Tui: carbonates barite

Casapalca: quartz calcite

Finlandia: quartz

Yankee Fork: intermediate volcanics ash-flow tuffs rhyolitic intrusions vein quartz

Western Cascades: volcanics altered volcanics altered intrisions

Golden Sunlight: clastlc seds. altered host breccias

quartz seiicite

Figure 6.13. Distributions of 6 180 in epither- m a 1 deposits.

amount of compositional s ca t t e r , they reveal a number of sys t ema t i c trends when examined on t h e basis of individual occurrences. T e large isotopic variability results f rom (I ) differing lk0 compositions of t h e host rocks, (2) differing ''0 compositions of t h e hydrothermal fluids, and (3) in teract ions between the host rocks and fluids over a range of low temperatures ( 1 0 0 ~ ~ t o 300 '~ ) and a t t a inmen t s of equilibrium t h a t produced hydrotherma minerals of variable and generally in termedia te "0 compositions.

As previously noted, most of t he common marine sedimentary ocks a r e isoto ically more variable and 8 enriched in ''0 (=LO t o 40 loo) than a r e the common igneous rocks ( '5 t o 12 O/oo; s e e Fig. 6.7). These petrologic distinctions accou for the obvious isotopic d i f ferences between t h e '0-enriched "unalteredQ' sedimentary host rocks of t he Geysers and Salton Sea geothermal areas and the Cor t ez and C lin epi thermal deposits, a s compared t o t h e less "0- enriched "unaltered" volcanic host rocks of nearly a l l o the r epi thermal deposits.

The derivation of nearly a l l geothermal and epi thermal fluids from meteo r i c sources of wa te r has also been mentioned reviously. This conclusion is based on t h e D and '6 compositions of these fluids, which although variable plot in close proximity t o or by la tera l "oxygen isotope shift" away from t e

1 6 me teo r i c water line (Fig. 6.8). The 6D and 6 0 compositions of hydrothermal fluids t h a t formed the epi thermal deposits of this discussion and those for many of their nearby present-day me teo r i c wa te r s a r e l isted in Table 6.6. Sources of da t a for most deposits of t he Basin and Range a r e from O'Neil and Silberman (19741, excep t for those of Carlin (Radtke e t al., 1980) and C o r t e z (Rye et al., 1974). Da ta for epi thermal deposits of other geographic locali t ies include Tolfa (Lombardi and Sheppard, 19771, Yankee Fork-Idaho Batholith (Criss and Taylor, 1983; Criss e t al., 19851, and others previously mentioned. On t h e basis of isotopic similarit ies t o pgesent-day me teo r i c waters and of widespread D and 0 depletions, t h e sources of most epi thermal fluids must have been local me teo r i c wa te r s (Table 6.6). Although magrnatic wa te r s may const i tu te a small proportion of these hydrothermal fluids, t hey have been de tec t ed with confidence f rom D analyses of fluid inclusions only a t Casapalca (Rye and Sawkins, 19741, for t h e ear ly and in termedia te s t ages of carbonate mineralization a t Creede (Bethke and Rye, 19791, and a t Golden Sunlight (Por ter and Ripley, 1985). Based on isotopic and other considerations, t h e fluids of Tui (Robinson, 19741, Finlandia (Kamilli and Ohmoto, 19771, and possibly the Comstock Lode (Taylor, 1973) may be par t ly of magmat i c origin. Thus, the compositions of hydrothermal fluids depicted by closed triangles on Figure 6.8, with the exception of those of Casapalca and Golden Sunlight, a r e i l lustrative of me teo r i c water-dominated epi thermal systems. Compositionally complex fluids such a s those of Creede (Bethke and Rye, 19791, Sunnyside (Casadevall and Ohmoto, 19771, Finlandia (Kamilli and Ohmoto, 1977) and pa r t s of severa l o ther hydrothermal sys tems a r e not por t rayed in Figure 6.8. Fluids f r o r 8 these deposits have exceedingly variable 6D and 6 0 values (Figs. 6.6 t o

6.8) and thus enclose large isotopic domains a s summarized by Taylor (1979). This variability is a t t r ibutable e i ther t o paragenetically distinct and (or) mixed sources of magmat i c and me teo r i c waters, a s deduced f rom t h e analyses of many samples, or t o subsequent contamination by deuterium-depleted ground waters t h a t were trapped in pseudosecondary inclusions (Foley e t al., 1982).

Distributions of t h e isotopic d a t a (Fig. 6.13) show t h a t a l t e red host rocks, regardless of igneous or sedimentary parentage, a r e variably deple ted in ''0 rela t ive t o thei r unaltered counterparts. hese t rends a r e evident f rom comparisons of t he "0 d a t a for unaltered and a l t e red host rocks a t Salton Sea-Cerro Pr ie to (Clayton e t al., 1968; Williams and Elders, 19841, Broadlands-Wairakei (Blattner, 1975; Clayton and Steiner, 19751, C o r t e z (Rye e t al., 19741, Carlin (Radtke et al., 19801, Tolfa (Taylor and Turi, 1976; Lombardi and Sheppard, 19771, Tonopah (Taylor, 1973), Bodie (OQNeil e t al., 1973), San Juan Mountains, Colorado (Taylor, 1974b; L e a e t al., 1984), Western Cascades, Oregon (Taylor, 1971; 1974a), and Golden Sunlight (Porter and Ripley, 1985). Although comparisons t o unaltered equivalents a r e lacking, a l t e red sedimentary host rocks a t t he Geysers (Sternfeld, 1981) and volcanic host rocks a t t h e Comstock Lode and Goldfield (Taylor, 1973; O'Neil and Silberman, 1974), and a t Yankee Fork (Criss and Tay r, 1983; Criss et al., 1985) a r e similarly depleted in "0 rela t ive t o normal magmat i c compositions. These ''0 depletions of volcanic and sedimentary host rocks a r e a corollary of t he "oxygen isotope shift" previously noted fo r many geothermal fluids (Fig. 6.8). They ar ise during geothermal-hydrothermal ac t ' i t y from isotope-exchange reactions between "0- depleted meteor i waters and "0-enriched host r o s and result in theC1'O enr ichment of t h e fluids and '$6 depletion of the rocks a s a consequence of ter-rock reactions. Progressive depletions of ''0 with increasing depth in host rocks of t he Cer ro Pr ie to (Williams and Elders, 19841, Geysers (Sternfeld, 1981, and references therein), and Salton Sea (Clayton e t al., 1968) a reas cannot be re la ted t o temperature-induced f rac t ionat ion trends, and thus mus t be caused by changes in fluid compositions with depth t h a t result f rom recharge and (or) water-rock reactions. Isotopic trends and permutat ions resulting from these phenomena range from subtle t o dramat ic , and they will be discussed extensively in t h e section t h a t follows.

Compositions of the fracture-controlled hydrothermal gangue minerals from geothermal sys tems and epi thermal deposits (Fig. 6.13) a r e largely compatible with isotopic e f f e c t s described in t h e foregoing discussion and from previous considerations of fractionation. The d a t a a r e f rom sources previously cited, except for t he "other" ca tegor ies of sediment- and volcanic-hosted deposits t h a t a r e from O'Neil and Silberman (1974; also s e e Table 6.6) and represent miscellaneous mines and prosp t s f rom t h e Basin and fS Range province. Values of 6 0 for these minerals generally occupy a range t h a t is in termedia te between t h e compositions of associa ted altered-unaltered host rocks (Fig. 6.13) and those of t h e geothermal or

Table 6.6 Compositions of 6D and 6180 in hydrothermal fluids of Tertiary epithermal deposits and some local meteoric waters

Epithermal F1 yQds Meteoric W 6 D 6 0

Basin and Range

Volcanic-Hosted Deposits

Bullfrog (BU) Aurora (A) Trade Dollar (TD) Wonder (W) Jarbidge (J) Rawhide (R ) Gilbert(G) Tonopah (T) Bodie (B) Cornstock Lode (CL)

Sediment-Hosted Deposits

Tenmile (TE) Humboldt (H) Cortez (CO) Carlin (CA) Manhattan (M)

Elsewhere

Volcanic-Hosted Deposits

Tolfa (TF)

Zoned Polymetallic Vein

Creede Sunnyside Tui (TU) Casapalca Finlandia Yankee Fork (Idaho Batholith, IB)

Western Cascades (WC) Golden Sunlight (GS)

complex complex

0 7

complex -1 6

*calculated from equation for the meteoric water line (eq. 12)

hydrothermal fluids (Table 6.6 and Fig. 6.8). The 6180 values of hydrothermal minerals formed by mineral- H 2 0 exchange reactions a r e determined by t h e t empera tu re of deposition, which controls t h e e x t e n t of isotopic fractionation, and fluid compositions, which in turn a r e controlled by t h e source of fluids, types of host rock, and water-host rock reactions. Isotopic e f f e c t s re la ted t o one or more of these determinants a r e qualitatively apparent f rom the d a t a of Figure 6.13. The t empera tu re control, with mineral-H20 f ract ionat ion increasing with decreasing t empera tu re ,

must account fo r t h e l a rge 1 8 0 enr ichments of siliceous hot-spring s in ters (22.2 and 23.6 O/oo) re la t ive t o vein qua r t z (mostly 3.9 to 12.3 O/oo) formed a t higher t empera tu res and depths in the Broadlands- Wairakei a reas (Clayton and Steiner, 1975; Blattner, 1975). Moreover, this t empera tu re contro of f rac t ionat ion is largely t h e cause of progressive j80 depletions of vein qua r t z and (or) ca l c i t e with increasing depth a t Broadlands (Blattner, 1975)) Cer ro P r i e to (Williams and Elders, 19841, Geysers (Sternfeld, 1981), Salton Sea (Clayton e t al., 1968)) and Wairakei

(Clayton and Ste iner , b975). However, t he parallel t rend of decreas ing d. 0 values with depth in host rocks of severa l geothermal areas , a s previously noted, cannot be a t t r ibu ted t o t empera tu re ef fects , but ins tead must resul t f rom changes in fluid composition. In addition, compositions of t h e gangue minera ls appear t o be influenced by those of thei r host rocks. For example , hydrothermal ca l c i t e and qua r t z associa ted with "0-enriched sedimentary host rocks of Carlin and C o r t e z a r e largely isotopiffly heavier than thei r counterpar ts i n relatively 0-deple ted volcanic host rocks of t h e o ther epi thermal deposits; particularly in t h e Basin and Range province where compositions of t h e me teo r i c waters a r e l eas t variable (Table 6.6). The isotopic composition of hydrothermal fluids, a s par t ly determined by source, may also leave i t s imprint on t h e minerals. Quar tz and ca l c i t e deposited f rom ''0-enriched magmat ic waters (Casapalca, Golden Sunlight, and possibly carbonates a t Creede) or f rom relatively undepleted me teo r i c wa te r s (Tolfa and Tui) a r e isotopically heavier t han thei r equivalents formed in hydrothermal sys tems having more deple ted fluids of me teo r i c origin such as those of t h e Basin and Range (see T le 6.6 and Fig. 6.13). The sequence of re la t ive % enr ichments among t h e vein minerals (q t z>ca l>bar>Kf) is generally consistent with t h a t obtained from experimentally derived f rac t ionat ion d a t a (Fig. 6.3). However, analyses for coexisting mineral pairs and t r ip le ts a r e conspicuously few, and they a r e largely suggestive of isotopic disequilibrium. This result is not surprising because of t h e relatively low t empera tu res (IOO°C t o 300°c), changes in fluid chemistry, and varied sequences of mineral paragenesis t h a t may cha rac te r i ze geothermal-epithermal systems, and which collectively render isotopic equilibrium unlikely. Investigations of fluid-mineral isotopic equilibria in geothermal sys tems by Clayton e t al. (19681, Bla t tner (19751, and Clayton and Steiner (1975) have demonstra ted t h a t qua r t z is most resistent t o isotopic exchange, whereas ca l c i t e and alkali feldspars may rapidly undergo re-equilibration and thus be susceptible t o compositional change during post- depositional s t ages of hydrother a l activity.

The principal t rends for '0 in host rocks and hydrothermal gangue minerals associated with epi thermal ac t iv i ty a r e summarized by the composite i l lustration given in Figure 6.14 (af ter Taylor, 1971; 1973). Volcanic rocks of t h e Western Cascades in Oregon host numerous zoned polymetall ic vein deposits of t h e base and precious metals. These deposits a r e mostly c lus tered within larger district-sized a reas of hydrothermal a l tera t ion (Field and Power, 1985) t h a t a r e cored by small granodiorite intrusions of Ter t iary age. According t o Taylor (19711, t h e volcanic country rocks (origin l y 5.5 t o 8 O/oo) a r e progressively ?d . . deple ted in 0 wlth Increasing proximity t o t h e intrusions (Fig. 6.14A; a f t e r Taylor, 19711, a s a consequence of hydrothermal a l tera t ion imposed by react ions between heated me teo r i c ground wa te r s ("-9 O/oo, Table 6.6) and the vol nic host rocks (now - 5 . 5 t o 5.5 O / o o . This trend of '0 depletion, which increases with intensity of a l tera t ion, is analogous t o t h e "oxygen isotope shift" of geothermal waters, but

opposite in direction. Volcanic host rocks of t he Tonopah dis t r ic t exhibit similar depletions of ''0 with progressive a l tera t ion (Fig. 6.148; a f t e r Taylor, 1973), and evidence for this having been a me teo r i c water- dominated hydrothermal sys te is additionally s t r ngthened by t h e pronounced 'gb depletions (= -7 t o I ~ ~ o o ) of associated quar tz , ca lc i te , and adular ia vein minerals. The a rea l distribution of 1 8 ~ - d e p l e t e d country rocks a r e large a t Bohemia in the Western Cascades (Taylor, 19711, Tonopah (Taylor, 19731, and a t Yankee Fork and other hydrothermally a l t e red a reas of t he Idaho Batholith (Criss and Taylor, 1983; Criss et al., 1985), and these "negative" isotopic anomalies form well-defined t a rge t s appropriate to the reconnaissance s t age of mineral exploration.

Waterxock ratios--In this sect ion we examine t h e sys temat ics of oxygen- and hydrogen-isotopic exchange between fluid and rock and present a n

B. X : OUARTZ :WHOLE ROCK O : CALCITE 0 : K FELDSPAR

VOLCANIC • • BRECCIAS

A. 4 . -

+8

+7

+6

+ 5 -

1 4

- t3-

g 2

. -

ODDlE • 0 . • RHYOLITE K

- . . - . ) * . - 8 .

* . ------------ s.ss./. .----------------------- . - I . I

:

LATE EUHEDRAL x "x x 7," QUARTZ .,@ ,,

MASSIVE

QUARTZ Ex 81100

QUARTZ IN FAULT BRECCIA x X

WEST END 5Y:F.. . ; . RHYOLITE

.

*. I . 8 0 ' ~ v a 1 u o of v o l c a n i c c o u n t r y rocks, Western C a s c a d e Range. O r e g o n

. . MIZPAH

o-. '''.,I;' . . . 8 : 5 5 TRACHYTE co

SANDGRASS - ANDESITE , , , , , , , , , , , , , , , , ,

-7 -6 - 5 .4 . 3 -2 - 1 0 1 2 3 4 5 6 7 8 9 10

8d8 (%.I

-3

-4

- 5

- e

Figure 6.14. Variations of 6 180 i n (A) volcanic host rocks adjacent to Ter t iary intrusions i n t h e W e s t e r n Cascades o f Oregon (a f te r Taylor, 1971, p. 7867, Fig. 8) and (B) volcanic host rocks and magmat ic and hydrothermal minerals i n the Tonopah d i s t r i c t of Nevada ( a f t e r Taylor, 1973, p. 755, Fig. 6 ) -

-

- z i 0 P m 0 . -

. I , , , , , / ! ! I I o 1 2 3 4 5 6 7 e 9 ( 0 1 r r 2 4 3 ~

D ~ s t o n c e from neorest ln t rus lve contact (an u n l l s of s tock d lometer )

application of these principles t o epi thermal systems. The application serves not only to impose const ra in ts on t h e hydrothermal environment, but also a s an i l lustration of the assumptions, problems, and uncer ta in t ies t h a t may be involved in modeling isotopic data. The development of t he principles of isotopic exchange in water-rock sys tems generally follows t h a t of Ohmoto and Rye (1974).

The final isotopic composition of water (6 a f t e r equilibration with rock is a function of: (a) t h e init ial (unexchanged) composition of t h e water (6 &) and rock (6: !, (b) t h e temperature of equilibration, which determines t h e fractionation f ac to r between rock and water (Ar-,), and (c) t h e ra t io of exchanged oxygen and hydrogen a t o m s in t h e water t o those in rock (wit-). This relationship, a f t e r Ohmoto and R y e (1974), is expressed a s

Epithermal precious-metal deposits a r e commonly hosted e i ther by volcanic rocks of i n t e rmed ia t e t o fe ls ic composition or by c l a s t i c or chemically precipi ta ted sedimentary rocks. Unaltered andesites, dacites, and rhyolites typically have 6180 and 6D values of about 7 '100 (Fig. 6.7) and -70 '100, respectively. Rocks of this compositional range contain approximately 50 weight-percent oxygen and a s much a s 0.11 weight-percent hydrogen (1 wt.-% H20), according t o t h e average analyses repor ted by Nockoids et al. (1978). Thus, t h e oxygen- and hydrogen-isotopic composition of a fluid t h a t has equilibrated with volcanic rock of these character is t ics can be determined from

and

where t h e coeff ic ients 1.8 and 100 represent ra t ios of t h e weight-percent oxygen in water (88.8%) t o t h a t in rock (50%) and the weight-percent hydrogen in water (1 1.2%) t o t h a t in rock (0.112%), respectively, and R is t he water:rock mass ratio. Because values of R represent proportions of water and rock t h a t have isotopically equilibrated, they a r e easier t o r e l a t e t o natura l sys tems than a r e values of t he a tomic ra t io w/r.

Isotopic compositions of sedimentary rocks a r e more variable than those of igneous rocks (Fig. 6.7). However, t he s t ra t igraphic sequences t h a t host epi thermal deposits typically consist of si l ty argil laceous limestone, dolomite, quar tz i te , and minor shale. If a "typical" host sequence contains subequal amounts of si l iciclastic and carbonate com wents, then i t would average approximately 16 '100 6 0 and

-60 '100 6D and contain about 50 weight-percent oxygen and perhaps 0.28 weight-percent hydrogen (2.5 wt-% H2.0). The final oxygen and hydrogen isotopic compos i t~on of a fluid t h a t has equilibrated with a sedimentary rock of this composition c a n be calcula ted f rom

and

The fractionation of oxygen isotopes between fluids and rock has been variously assumed t o be similar t o those of smect i te-H 0 (Cathles, 1983), plagioclase feldspar (An 0)-H20 ( f ay lo r , 1974a, 1979; Ohmoto and Rye, 1972; Green e t al., 1983) and muscovite-H20 (Spooner e t al., 1977). The assumptions a r e based on comparisons of these mineral-H20 fractionations with exper imenta l rock- H 2 0 fractionations and with t h e oxygen-isotopic compositions of naturally a l t e red rocks and secondary minerals. For t h e i l lustrations t h a t follow, we have used fractionation factors derived from the ~ l a ~ i o c l a s e feldspar (An30)-H20 equation of OtNeil andu Taylor (1967)

I000 I n a = 2.68 ( 1 0 6 / ~ * ) - 3.29 (18 )

Use of this equation r a the r than the more r ecen t one of Matsuhisa e t al., 1979 (from eqs. 7 and 10 in Table 6.3) was done t o maintain continuity between our resul ts and those of other investigators (see Taylor, 1979, and above). These f rac t ionat ions a r e in termedia te between those derived f rom the muscovite-H20 and smectite-H 0 curves over the t empera tu re range from 1 0 0 ~ 8 to 3 0 0 ~ ~ . The f rac t ionat ion of hydrogen isotopes between fluids and rock is similarly assumed t o be equivalent t o t h a t of bioti te or chlorite-H20 (Taylor, 1974a). Our computations a r e based on chlorite-H 0 fract ionat ion f ac to r s which were taken from Taylor f1979, Fig. 6.2).

Using t h e equations, rock compositions, and mineral-H20 systems described above, t h final

78 1 - isotopic composition of me teo r i c water ( 6 - -1 '/PO and 6% = -120 .O/OO) and magmat i c waters ( Jgq: = 7.3 '/oo and 6% = -60 '/oo) were computed a t t empera tu res of 1 0 0 ~ ~ to 300°C and R values of 0.01 t o 10. The init ial isotopic composition of t h e me teo r i c water is typical of present-day me teo r i c wa te r in western Nevada (Table 6.6) and t h e se lec ted t empera tu res and R values a r e considered re levant t o epi thermal systems. The results of t h e com putations, portrayed in Figure 6.15, i l lus t ra te the sys t ema t i c s of isotopic exchange between me teo r i c or magmat i c waters and rocks in the Basin and ange province. Fi rs t , me teo r i c water is enriched in '0 and D under most geologic conditions, whereas magmat i c wa te r is enriched in deuterium but may be e i ther deple ted

in ''0 thro h exchange with igneous rocks or enriched in "0 via exchange with sedimentary rocks. Second, the magnitude of t h e isotopic enrichment or depletion varies inversely with t h e water:rock mass ratio. Third, the change in fluid 6180 values exceeds t h a t of 6D values a t high water:rock mass ratios (?,I) but t h e reverse is t r u e a t low ra t ios ($1). This e f f e c t results from t h e small H:O mass r a t io in rocks re la t ive t o t h a t in water and f rom the generally larger f rac t ionat ion f ac to r s for hydrogen than for oxygen. Accordingly, t he isotopic influence of rock hydrogen is significant only a t low water:rock mass ratios, but is potentially larger than t h a t of rock oxygen. A c ollary t o this e f f e c t is t h a t t h e generally small '"0 shiftf' observed for geothermal waters (Fig. 6.8) must be a product of isotopic exchange a t relatively high water:rock mass ratios. Fourth, a t equivalent t empera tu res and wa te rxock mass ratios, fluids t h a t have equilibrated with sedimentary rock a r e isotopically heavier than those t h a t have equilibrated wi igneous rock because t h e former is enriched in"0 and D and contains a g rea te r quantity of hydrogen re la t ive t o t h e la t ter .

The calcula ted and analyzed isotopic compositions of t h e hydrothermal fluids associated with epi thermal deposits f a l l in the range -15 t o 5 O/oo 6180 and -150 t o -90 O/oo 6D (Table 6.6 and Fig. 6.8). Consequently, these fluids a r e isotopically deple ted re la t ive t o magmat i c wa te r or evolved magmat i c water , but a r e similar t o me teo r i c water and i t s evolved counterpar ts (Fig. 6.15). Because i t is unlikely t h a t magmat ic fluid will be isotopically depleted in deuterium through equilibration with rock (most mineral-H20 f ract ionat ion f ac to r s a r e negative fo r hydrogen), i t is concluded t h a t me teo r i c wa te r is t h e predominant fluid in epi thermal systems. However, if fluid mixing is commonplace, then a minor component of magmat ic water cannot be precluded on t h e basis of isotopic considerations.

The isotopic composition of a convected fluid may di f fer f rom t h a t of a s t a t i c fluid because, among other factors, the convected fluid has equilibrated over a range of t empera tu res and water:rock mass ratios. Therefore, t o model the evolution of western Nevada me teo r i c wa te r during convection, t h e isotopic composition was calcula ted a t 20°C intervals f rom l0fI0C t o 300°C. The final composition of t h e water (6,) computed a t e a c h t empera tu re became the init ial composition of t h e water (6;) for t h e calculations a t t he next higher temperature . Since wa te rxock mass ra t ios may vary according t o t h e convection path, t h e ser ies of calculations were performed a t R values of 0.01, 0.1, 1, and 10.

The results a r e presented in Figure 6.16 and compared t o compositions of t h e fluids responsible fo r those volcanic-hosted (Fig. 6.16A) and sediment-hosted (Fig. 6.16B) epi thermal deposits l isted in Table 6.6. The computations imply tha t co vect ing me teo r i c wa te r i s progressively enr iched in "0 and D during heat ing and exchange with unaltered wall rock. This wa te r will be isotopically heavier t han non-convecting me teo r i c water (e.g. pore fluid), under identical conditions of equilibration, because of i t s "history" of exchange (compare Figs. 6.15 and 6.16). Moreover,

Magmatic 3000~ Water

Figure 6.15. Variations of 6~ and 6 180 in fluids that equilibrate with (A) volcanic and (B). sedimentary rock as a function of the initial fluid composition, temperature, and water: rock ratio.

isotopically light me teo r i c water may a t t a in compositions similar t o those of magmat ic or evolved magmat ic wa te r through exchange a t low water:rock mass ra t ios (L0.1). Therefore, t he distinction between a magmat i c or me teo r i c source of fluids and t h e demonstration of mixing between magmat ic and me teo r i c waters cannot be made solely on the basis of t he ca lcula ted composition of a hydrothermal fluid.

The oxygen-isotopic composition of meteor ic wa te r in t h e Basin and Range province has remained essentially unchanged since the Early Tertiary, whereas t h e hydrogen-isoto i c composition may have g decreased slightly (10-20 loo) in response t o the c l imat ic cooling (Sheppard e t al., 1969; Taylor, 1973). Thus, t he isotopic composition of me teo r i c water in this region may be regarded a s a n approximation of t h e init ial composition of Ter t iary me teo r i c hydrothermal fluids. Meteor ic wa te r compositions near t h e Aurora, Rawhide, Gilbert , Tenmile, Humboldt, C o r t e z and Manhattan epi thermal deposits have 6D values between -130 and -120 O/oo (Table 6.6). Accordingly, t he isotopic compositions of t h e hydrothermal fluids for

these deposits can b e in terpre ted in t e rms of t h e calculated curves displayed in Figure 6.16. Isotopic da ta for t he remaining deposits in the Basin and Range province (Table 6.6) a r e more appropriately compared t o ca lcula ted curves t h a t have been shifted, re la t ive t o those in Figure 6.16, in t h e direction and t o t h e ex ten t t h a t t he init ial compositions of t he fluids differed f rom -16 0 / o o 6 ~ ~ 0 and -120 O/oo 6D. Assuming t h a t temperatures of water-rock equilibration were typically between 150°C and 3 0 0 ' ~ (hachured fields in Fig. 6.16), a s indicated by fluid-inclusion studies, then most of t h e epi thermal sys tems were character ized by high water:rock mass ratios (>0.5). However, t he model calculations imply t h a t t h e Bodie, Tonopah, and Tenmile fluids evolved a t unusually low water:rock mass ratios (W.01-0.2) and t empera tu res (510O0C). Furthermore, t h e model cannot account for t h e apparent decrease in t h e 6D composition of the Carlin and Comstock Lode fluids.

Most of t he water:rock ratios determined above should be regarded a s t en ta t ive (particularly t h e anomalously low values) because processes o ther than fluid-rock exchange may have influenced t h e isotopic

Figure 6.16. Variations of CD and 6180 in fluids during convection through (A) volcanic and (B) sedimentary rock as a function of temperature and waterzrock ratio.

composition of the fluids. Fo r example, boiling in t h e convectio sys tem would enr ich t h e evolved meteor ic fluid in ''0 and D along a t rend subparallel t o t h e MWL (Truesdell et al., 1977; s e e Rad tke et al., 1980). Unexchanged me teo r i c wa te r may become involved a t t h e s i t e of mineralization through mixing or with t h e complete equilibration (alteration) of t he host rocks. This e f f e c t ould drive t h e fluid composition towards

15 the init ial 6 0 and dD values. Inaccurate water:rock ratios may also resul t if t h e isotopic composition of modern me teo r i c wa te r is used a s an approximation of Ter t iary me teo r i c wa te r in a reas t h a t have undergone considerable uplift o r subsidence. Similarly, 6 D analyses of ex t r ac t ed inclusion fluids t h a t contained a significant f rac t ion of secondary fluids could lead t o erroneous conclusions (see Foley e t al., 1982).

The final i sotopic composition of t h e rock (6 f) t h a t equilibrated with t h e me teo r i c fluid a t e a c h interval in t h e convection model was ca lcula ted f rom t h e relationship

and por t rayed a s curves in Figure 6.17. These results ind' a t e t h a t rocks become progressively depleted in '$0 and D with a n in e a s e in t h e water:rock mass ratio+ The depletion in "0 is negligible a t low ra t ios because of t h e overwhelming abundance of rock oxygen. With increasing temperature , both the 6180 and 6D values increase a t water:rock mass r t ios of 1

1'8 or less, whereas a t a r a t io of 10 the 6 0 value decreases and t h e 6D value increases.

Decreases in t h e 6180 values of hydrothermally a l t e red rocks nea r cen te r s of hydrothermal ac t iv i ty have been documented for t h e Tonopah epi thermal Au-Ag deposit (Taylor, 19731, zoned polymetall ic veins of t h e Bohemia dis t r ic t in the Western Cascades (Taylor, 19711, Yankee Fork and other a l t e red and mineralized locali t ies of t he Idaho Batholith (Criss and Taylor, 1983; Cr iss e t al., 1985), and several volcanogenic massive sulfide depo i t s (see Franklin e t

1 2 al., 1981). The whole rock 6 0 compositions a t Tonopah range f rom values typical of unaltered igneous rocks (5.5 t o 10.0 O/oo) t o nearly -6 permil and whole rock 6D compositions range from -150 t o -135 permil. The field represent72 by six Tonopah samples for which both 6D and 6 0 analyses have been repor ted is shown in Figure 6.17. Compositions of these samples a r e most consistent with isotopic exchange a t high water:rock mass ratios (> I ) , although t h e 6 D values a r e somewhat larger than predicted. On t h e basis of whole-rock 1 8 0 depletions observed a t Tonopah and elsewhere, water:rock mass ratios fo r many epi thermal d is t r ic ts throughout t h e western U.S. range f rom about 0.2 t o 2 (Taylor, 1974a) and these ra t ios a r e broadly similar t o those of geothermal sys tems t h a t range from about 0.15 a t t he Geysers (Sternfeld, 1981) through 0.45 and 1.3 a t Salton Sea (Clayton et al., 1968) and C e r r o Pr ie to (Williams and Elders, 19841, t o a s large a s 4.3 a t Wair e i (Clayton and Steiner, 1975). The ou te r haloes of % depletion extend beyond t h e megascopically identifiable e f f e c t s of a l tera t ion and anomalies of t r a c e and minor e l emen t s (Taylor, 1971; Green e t al., 1983).

y ~ . I SUMMARY

Figure 6.17. Variations of 6~ and 6180 in (A) volcanic and (B) sedimentary rocks that have equilibrated with convected meteoric water over a range of temperatures and water :rock ratios.

Accordingly, these cones of 1 8 0 depletion may be potentially useful in t h e exploration for hidden mineral deposits (see Criss and Taylor, 1983; Criss e t al., 1985).

Two main conclusions pertaining t o epi thermal sys tems may be summarized f rom this discussion on the exchange of oxygen and hydrogen isotopes between fluids and rock. First , meteor ic wa te r is t h e predominant, and possibly only, source of fluid in most of t h e epi thermal gold-silver deposits studied t o date. Second, a t reasonable t empera tu res of 1 5 0 ' ~ t o 300°c, t h e isotopic d a t a for hydrothermal fluids and thei r associa ted a l t e red rocks both imply t h a t high water:rock mass ratios prevailed during mineral deposition. Because typical porosities l imi t bulk water:rock mass ratios to generally less than 0.1 (e.g. 10% porosity is equivalent t o a r a t io of 0.04), then fluid-rock equilibration and mineral deposition must have occurred in open sys tems through which masses of fluid c i rcula ted t h a t were equivalent t o or larger than those of the rocks.

This overview of light-stable isotopes in the epi thermal environment is prefaced by a review of t h e general principles of equilibrium isotope-exchange fractionation, including t h e equations and graphical portrayal of f rac t ionat ion e f f ec t s be twee com on isotop' compounds, and a summary of D,73C, '0, and "S distributions in geologically important habitats. The available d a t a for these isotopes in geothermal and epi thermal sys tems a r e largely consistent with trends t h a t might be inferred from experimental-theoretical considerations and known distributions in common rock types and mineral deposits. Nonetheless, t he epi thermal d a t a exhibit isotopic character is t ics t h a t a r e in termedia te between those of near-surface geothermal sys tems a d deeper hydrothermal deposits. Depletions of D and ' 0 in t h e host rocks, minerals, and inclusion fluids of the epi thermal deposits, and compositional similarit ies of these fluids t o present-day me teo r i c waters, suggest t h a t t h e hydrothermal fluids were predominantly ground waters of me teo r i c origin; although several deposits (Casapalca, Golden Sunlight, and possibly Creede) may have had varying amounts of a magmat ic component. Sources of carbon in t h e carbonate minerals a r e not uniquely defined, and probably have originated f rom nearby sedimentary host rocks (Carlin and ~ o r t e z ) , magmas, and possibly from biogenic and other provenances. Those of sulfur a r e considered t o have been derived f rom magmas or igneous rocks, sul fa tes ( S u n n ~ s i d e and Tolfa) and sulfides (Carlin, Cor tez , and Guanajuato) of sedimentary rocks, and s e a water (Pueblo Viejo and Tui). Temperature-controlled f rac t ionat ions over l a t e ra l or ver t ica l gradients ranging from 300°C t o IOOOC outward and upward in hydrothermal systems should result in heavy-isotope

r ichments of about 12 O 1 3 c for ca lc i te , 14 '100 "0 for quartz, 12 Oioo for ca lc i te , 26 O/oo 3 4 ~ for sulfates, and 1.7 O/oo 3 4 ~ for pyrite. Although these "hypothetical" f rac t ionat ion t rends have not been reported f rom any of t h e epi therm 1 depos' ts studied t o date , they have been noted for ''C and "0 in ca l c i t e and quar tz of geothermal sys tems (Broadlands, Ce r ro Prieto, Geysers, Salton Sea, and Wairakei). Their absence f rom epi thermal deposits may possibly be the result of inadequate sample representation, or of isotopic changes in fluid composition caused by boiling, redox and (or) water- rock reactions, disequilibrium, and di f ferent sources of t h e isotopic elements. Near-surface occurrences of supergene sulfates, which have formed by oxidation of ascending H2S or of preexisting lfides in the host rocks, a r e markedly deple ted in "S re la t ive t o t h e deeper hypogene sulfates. The a l t e red sedimentary and volcanic host rocks of both epi thermal deposits and geothermal sys tems a r e conspicuously depleted in 1 8 0 re la t ive t o lbheir peripheral and unaltered equivalents. These 0 depletions a r e a corollary of t h e well-known "oxygen isotope shift" of geothermal fluids, and they form in hydrothermal sys tems character ized by relatively high wa r:rock mass ra t ios (21). The resul tant "negative" IgO anomalies may serve a s a useful guide t o mineral exploration because they a r e isotopically unambiguous and areal ly


extensive. Isotopic investigations of epi thermal deposits should be continued, especially in conjunction with geologic and o the r topical studies, and particular emphasis should be given t o sys t ema t i c and three- dimensional sampling of host rocks beyond ore-bearing structures.

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CHAPTER 6

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Chapter 7 GEOLOGIC, MINERALOGIC, AND GEOCHEMICAL CHARACTERISTICS

OF VOLCANIC-HOSTED EPITHERMAL PRECIOUS-METAL DEPOSITS Daniel 0. Hayba, Philip M. Bethke Pamela Heald, and Nora K. Foley

INTRODUCTION

In Chapter I, R. W. Henley summarized our understanding of the chemical and hydrodynamic s t ruc tu re and the t ranspor t properties of ac t ive hydrothermal systems, with particular emphasis on ter res t r ia l magmatic-hydrothermal systems. Such an overview is especially valuable because act ive geothermal systems a r e modern "archetypes" of t he ancient systems which concentra ted me ta l s in their upper portions t o form epi thermal ore deposits. More than any other fac tor , t h e study of ac t ive sys tems has provided the framework on which the observations on epi thermal deposits have been arranged in t h e relatively recent development of comprehensive models of epi thermal ore formation. The Principle of Uniformitarianism has served us well in th is instance.

In this chapter, we focus on observations on epithermal ore deposits in continental silicic t o andesit ic volcanic terranes. Volcanic-hosted deposits offer t he most d i rec t comparison with many of the well-studied modern geothermal systems. We f i rs t compare the a t t r ibutes f rom a number of epi thermal o re deposits and show how they may be used t o identify two important, and distinct volcanic-related hydrothermal environments. We then examine t h e best-studied deposit of each type: Creede and Summitville, both of which a r e located in t h e San Juan Mountains in southwest Colorado. In so doing, we a r e able t o examine epi thermal deposits for evidences of processes t h a t a r e now occurring in geothermal systems. Finally, we use t h e observational base and in terpre ta t ions derived from each deposit t ype t o develop generalized "geothermal" models of mineralization. The models have been taken, in large par t , from t h e excel lent synthesis by Henley and Ellis (1983). We feel t h a t thei r models a r e soundly based on a myriad of d i rec t observations on ac t ive geothermal systems, and a r e consistent, for t h e most par t , with t h e observations made a t Creede and Summitville.

SUMMARY OF THE CHARACTERISTICS O F VOLCANIC-HOSTED EPITHERMAL ORE DEPOSITS

In 1981, Buchanan published a valuable compilation of se lec ted observations f rom over 60 gold-silver vein deposits in unmetamorphosed volcanic- to-subvolcanic environments. These d a t a and t h e in tegra ted model derived f rom them have formed a useful basis for numerous subsequent analyses of t h e character is t ics of epithermal deposits, some of which include Giles and Nelson (19821, Ashley (19821, Bonham

and Giles (1983), Heald-Wetlaufer e t al. (19831, Berger and Eimon (19831, Sil l i toe and Bonham (1984), Heald e t al. (19861, and Bonham (1986). Because of the large number of deposits in Buchanan's 1981 study, a deta i led evaluation of the d a t a base was not possible. Heald e t al. (1986) undertook such a deta i led study of 16 carefully se lec ted, well-studied, Ter t iary volcanic- hosted epi thermal deposits. Using this d a t a base, Heald e t al. (1986) showed t h a t two types of epithermal deposits could be distinguished and t h a t Buchanan's d a t a base supplemented and supported this conclusion. The t w o types, distinguished primarily on the basis of vein and a l t e ra t ion mineralogies, a r e the Adularia-Sericite t ype and t h e Acid-Sulfate type. The character is t ic f ea tu res of t hese two main types a r e shown in Table 7.1 and a r e discussed below. Adularia- Sericite-type deposits a r e f a r more numerous than the Acid-Sulfate-type deposits (Table 7.21, and so predictably, t he re is more information on t h e former.

Major telluride deposits (e.g., Cripple Creek, Colorado) were excluded in t h e study by Heald e t al. (1986) because thei r unique mineral assemblage suggests t h a t t hese deposits make up a distinct class (Heald-Wetlaufer et al., 1983; Bonham and Giles, 1983; Bonham, 1986). Hot-spring-type deposits (see Silberman and Berger, 1985, this volume; Berger and Silberman, 1985, th is volume) were not well represented in thei r compilation because of a paucity of published descriptive data.

Character is t ics of Adularia-Sericite-type Deposits

St ructura l setting--The most common regional s t ructura l s e t t i ng fo r Adularia-Sericite-type deposits is along the margins of calderas, although other t ec ton ic se t t ings (typically structurally complex volcanic environments) a r e not uncommon. The importance of t h e ca ldera se t t i ng lies in t h e excel lent plumbing system i t provides for hydrothermal circulation (Lipman et al., 1976; Steven and Lipman, 1976). I t is important t o note, however, t h a t relatively f ew calderas in t h e western United S ta t e s have been mineralized (McKee, 1979; Rytuba, 19811, and in t h e San Juan Mountains, Colorado, only about 113 of t h e known calderas have had significant mineral production (Steven and Lipman, 1976).

Size of deposit--There i s a large range in t h e s ize of t h e Adularia-Sericite-type deposits. Guanajuato, a silver- and base-metal-rich dis t r ic t , covers a surface a rea of approximately 190 s q km. Districts with relatively low base-metal contents , such a s Oatman, tend t o be considerably smaller (12 sq km); Comstock is an important exception. The length-to-width r a t io

Table 7.1--Characteristics of the adularia-sericite type and acid-sulfate type deposits (compiled from Heald et al., 1986).

Acid-Sulfate Adularia-Sericite

Structural setting Intrusive centers, 4 out of the 5 studied related to the margins of calderas

Structurally complex volcanic environments, commonly in calderas

Size 1ength:width ratio

relatively small equidimensional

variable; some very large usually 3:l or greater

rhyodacite typical silicic to intermediate volcanics

Host rocks

Timing of ore and host

similar ages of host and ore (<0.5 m.y.)

ages of host and ore distinct ( > I m.y.)

Mineralogy enargite, pyrite, native gold, electrum, and base- metal sulfides

Chlorite rare no selenides Mn-minerals rare sometimes bismuthinite

argentite, tetrahedrite, tennantite, native silver and gold, and base-metal sulfides

chlorite common selenides present Mn gangue present no bismuthinite

Production data Both gold- and silver- rich deposits

noteworthy Cu production production

Both gold- and silver- rich deposits

variable base-metal

Alteration Advanced argillic to argillic (* - sericitic)

Sericitic to argillic

Extensive hypogene alunite Major hypogene kaolinite No adularia

supergene alunite occasional kaolinite Abundant adularia

Temperature

Salinity

Source of fluids

1 to 24 wt% NaCl eq.3 0 to 13 wt% NaCl eq.

Dominantly meteoric, possibly significant magmatic component

Dominantly meteoric

Source of sulfide sulfur

Deep-seated, probably magmatic

Deep-seated, probably derived by leaching wallrocks deep in system

Source of lead Volcanic rocks or magmatic fluids

Precambrian or Phanerozoic rocks under volcanics

'could be secondary in some districts. 2~imited data, possibly unrelated to ore. 3~alinities of 5 to 24 wt% NaCl eq. are probably related to the intense acid-sulfate alteration which preceded ore deposition.

of t h e surface projection of mineralized veins in th is t ype of deposit is generally on t h e order of 3:l. The ver t ica l range of mineralization is usually on t h e order of 400 t o 700 me te r s compared t o s t r ike lengths of severa l kilometers.

Host rocks--The composition of the host rocks of Adularia-Sericite-type deposits ranges from rhyolit ic t o andesit ic, and o re i s commonly hosted by severa l d i f ferent compositional units within a district . In a few of t h e districts, t he ore fluids also mineralized associated sediments (Creede, Guanajuato) o r intrusive rocks (Silver City, Idaho), but o re is mainly confined t o t h e volcanic rocks. The occurrence of o re in severa l lithologies in Adularia-Sericite deposits implies t h a t composition of t h e host rock(s) is not a controll ing fac tor . The lack of a genet ic t i e t o the host is fur ther implied by t h e f a c t t h a t ore deposition in Adularia- Sericite-type deposits almost always occurred more than 1 m.y. subsequent t o the formation of t h e host rock.

Mineralogy--The mineralogy of the Adularia- Sericite-type deposit is character ized by t h e presence of vein adularia and ser ic i te and by the absence of both hypogene a luni te and the assemblage: enargi te + pyr i te +/- covelli te (Table 7.2). Chlorite is also character is t ica l ly present. Although alunite is found in some of these deposits, in each case i t appears t o be a near-surface supergene occurrence, unrelated t o t h e primary ore-forming system.

Metal ratios--The high silver-to-gold production ra t ios of most of t he Adularia-Sericite-type deposits r e f l ec t t h e abundance of native silver, silver sulfides, and sulfosalts. Districts such a s Round Mountain, Nevada, and Oatman, Arizona, have low silver-to-gold ratios. The precious meta ls in these two dis t r ic ts a r e present mainly a s nat ive gold, native silver, and e lec t rum; silver sulfides and silver sulfosalts a r e r a r e (Table 7.1). Base-metal production is also usually low for t h e more gold-rich deposits. In most cases th is r e f l ec t s t h e lack of base-metal sulfides and sulfosalts, but base-metals a r e somet imes not recovered in milling operations and therefore production d a t a may not always accurate ly ref lec t t h e o re mineralogy. The fairly large sample s ize of t h e Adularia-Sericite-type deposits in Table 7.2 permits speculation t h a t t h e r e may be a continuum from the base-metal-rich, silver- r ich districts, such a s Colqui, Peru t o t h e base-metal- poor, gold-rich districts, such a s Oatman.

Wallrock alteration--The a l tera t ion pa t t e rns of Adularia-Sericite-type deposits a r e not yet well defined, due in par t t o the lack of detailed a l tera t ion studies. In general, Adularia-Sericite-type deposits a r e cha rac te r i zed by the predominance of s e r i c i t i c alteration* t h a t of ten borders a silicified zone near the vein. Also near the vein, fine-grained potassium feldspar and/or chlor i te a r e of ten disseminated in t h e wallrock. The ser ic i t ic zone typically grades outward in to a propylit ic zone. An argil l ic zone between t h e ser ic i t ic and propylitic zones is somet imes present. A t Creede, Pachuca (Mexico) and Oatman, t h e a l tera t ion over the o re body has been described a s a ser ic i t ic "cap", in terpre ted t o be t h e result of t h e condensation of acid volati les released a t depth during boiling (Barton et al., 1977; Buchanan, 1981). In many, if not most districts, t he outermost propylitic a l tera t ion

zone (typical of both Acid-Sulfate- and Adularia- Sericite-type deposits) appears t o have formed prior t o o re deposition and may be unrelated t o t h e ore- forming hydrothermal system.

*The t e rm "sericitic" is used in this chapter in the sense of a l tera t ion consisting of a mica-type mineral (e..g., i l l i te) + quar t z + pyrite, including mixed-layer i l l~ t e - smec t i t e in which i l l i te layers a r e predominant.

Thermal histor --For the Adularia-Sericite-type d e p o s ~ t s , iluid-inclu5on studies t ied t o a deta i led pa;agenktic sequence (compiled in Hayba, 1983) show t h a t most ore deposition occurred a t t empera tu res between 2 0 0 ' ~ and 300°C, with late-stage fluids typically depositing only gangue minerals between 140°C and 200°C. Of the 11 Adularia-Sericite deposits evaluated in Heald e t al. (19861, boiling has unequivocally been demonstra ted t o be associa ted with precious-metal deposition only a t Colqui (Kamilli and Ohmoto, 1977). Precious-metal deposition due t o boiling has been proposed for o ther d is t r ic ts based on less definit ive evidence (cf. Guanajuato (Buchanan, 1979), Tonopah (Fahley, 1981)). Roedder (1984, p. 426) notes t h a t some of t h e evidence for boiling a t Guanajuato and Tonopah is ambiguous and t h a t t h e occurrence of boiling in these deposits rnay have been overs ta ted . Boiling also has been noted a t Eureka (Casadevall and Ohmoto, 19771, Pachuca (Drier, 19761, and Creede (Barton e t al., 1977) during t imes other than precious-metal deposition. Although boiling is an e f f ec t ive mechanism for depositing o res (Henley e t al., 1984; Henley, 1985, this volume; Drummond and Ohmoto, 1985; Reed and Spycher, 1985, this volume), i t is not t h e only mechanism. In t h e study of any o re deposit, t h e evidence for boiling should be crit ically evaluated before assuming i t contributed t o precious- me ta l deposition. Bodnar e t al. (1985, this volume) discuss problems in using fluid-inclusion evidence t o document boiling in ore deposits. In addition t o boiling, t he mixing of fluids from two o r more sources has played a role in t h e thermal history of Adularia- Sericite-type deposits a s has been documented by fluid-inclusion and stable-isotope studies a t Creede (discussed in detail later) .

Composition of fluids-Salinit ies, deterrnined f rom freezing point depression measurements, range f rom 0 t o 13 wt.-% NaCl equivalent for fluid inclusions f rom those Adularia-Sericite-type deposits evaluated. Most of t he deposits have consistently low salinit ies, usually less than 3 wt.-% (Hayba, 19831, but Creede and Colqui have unusually high salinity fluids, ranging f rom 5 t o 12 wt.-% NaCl equivalent (Woods et al., 1982; Kamilli and Ohmoto, 1977). The high salinit ies in these two deposits may have a bearing on thei r relatively high base-metal contents because chloride complexes a r e an e f f ec t ive means of transporting base me ta l s (Barnes, 1979; Henley, 1985, this volume). On t h e other hand, fluid inclusions f rom the Sunnyside (nine, Eureka mining district , Colorado, which is also very base-metal rich, have low salinit ies (0-3.6 percent). Hedenquist and Henley (1985) and Henley

132 CHAPTER 7

Table 7.2--Selected mineralogical and production data for 54 epithermal districts (from Heald et al., 1986). Both types of districts are ordered from higher to lower base-metal production. The last 14 districts may represent a distinctively low base-metal subtype of the adularia-sericite-type district. AD, adularia; AL, alunite; SS, sulfosalts; AGS, silver sulfides; SP, sphalerite; GN, galena; EN, enargite; BA, barite; RC, rhodochrosite; EL - fluorite; X - denotes presence.

District

Base ~etal'

Ag : Au percent AD AL SS AGS SP GN EN BA RC FL

ACID-SULFATE TYPE

*Red Mtn., Colo. 68 high X X X X X X X X X X *Julcani, Peru 467 mod X X X X X X X *Lake City 11, Colo. 5 23 mod X X X X X X X *Summitville, Colo. 1.2 5 ? X X X X X *Goldfield, Nev. 0.3 1 X X X X X

ADULARIA-SERICITE TYPE

*Lake City (12, Colo. 5

*Colqui, Peru Tovar, Mex. Parral, Mex. Bohemia, Oreg. Namiquipa,

Me 7. *Eureka, Colo. *Creede, Colo. "Guanajuato, Mex. Guanacevi, Mex. Yoquivo, Mex. Calico, Calif. Fresnillo, Mex. *Pachuca, Mex. El Tigre, Mex. Great Barrier, New Zea. Silver Peak, Nev. Piz Piz, Nicaragua, "Tonopah, Nev. Zalcualpan, Mex. Tayoltita, Mex. Temalscaltepec, Mex. Golden Plateau, Aus. Mohave, Calif. Divide, Nev. Bodie, Calif. Stateline, Utah Guadalupe, Mex. Mogollon, N. Mex. *Corns Lock, Nev. *Silver City, Idaho *De Lamar, Idaho Gold Circle, Nev. National, Nev. Aurora, Nev. Ramsey-Tala, Nev. Republic, Wash. Searchlight, Nev. Rawhide, Nev. Ocampo, Mex.

259 2 6 45 150 6

Ag only 47

400 200

100-500 7 4

1200 64

200 162

4-30 24 3 2

80-110 Ag only

5 1

high 17.5

10 4-20

9 8.5 7.5 5

0-10 6-12 mod mod 4

3.5 3 3

1-4 2 2

1.4 1

1 ow minor minor v low v low v low rare

1 X X X X X X 1 X X X X X 1 X X X X 1 X X X X 1 X X X 1 X X 1 X X X 0.3 X X 0.2 X X X 0.03 X X X X

X X X X X X X X X X X X X X X X

X X X

Table 7.2--Selected mineralogical and production data for 54 epithermal districts (from Heald et al., 1986)--(continued)

District

Base Metal 1

Ag : Au percent AD AL SS AGS SP GN EN BA RC FL

ADULARIA-SERICITE TYPE (continued)

Rochester, Nev. Tuscarora, Nev.

*Round Mtn. Nev. Cornucopia, Nev. Wonder, Nev. Seven Troughs, Nev. El Oro, Mex. Jarbidge, Nev. Gilbert, Nev. Hayden Hill, Calif. Katherine, Ar iz.

*Oatman, Ariz.

113 44-100

0.2 6 8 94 5.4 7 3

Au only 1.5 3

0.4

*Denotes the 16 districts studied in detail in Heald eL al., 1986; the others were modified from Buchanan (1981).

'~ase-metal production, usually as sulfides, rarely as metals, as a percent of the total tonnage (Buchanan, 1981).

2~lunite younger than mineralization.

3~lunite shallow only, may be supergene.

4~lunite older than mineraliza~ion.

5~lack (1980) grouped the Galenea (Henson Creek) and Lake (Lake San Cristobal) districts, Colo. into the Lake City district; the earlier ores (Lake City I ) occur primarily in the Galena district and the later ones (Lake City 11) in the Lake dis~rict.

6 ~ a ~ a for the Finlandia vein.

7 ~ a t a mainly for the Sunnyside vein.

(1985, this volume) have noted t h a t t h e in terpre ta t ion of salinity from measurements of fluid inclusions in some deposits can be skewed t o higher values if t h e fluids have relatively high gas contents, because C 0 2 and H2S contr ibute t o t h e f reezing point depression. However, t he presence of C 0 2 in these epi thermal deposits has not been documented. Although high concentrations of C 0 2 can easily be de tec t ed by t h e crushing t e s t techniques described by Roedder (1970), Bodnar e t al. (1985, this volume) show t h a t inclusions trapped under "epithermal" conditions may contain significant amounts of C 0 2 which go undetected. A few analyses of inclusion fluids from Creede (Roedder, compiled in Hayba, 19831, Eureka (Casadevall and Ohmoto, 19771, and Colqui (Kamilli and Ohmoto, 1977) indicate C 0 2 concentrations of less than 0.5 molal in a lmost all samples.

Paleodepth--Paleodepth e s t ima tes from both geologic reconstructions and pressures es t imated from fluid-inclusion studies show t h a t most of t he Adularia- Sericite-type deposits appear t o have formed a t paleodepths of 300-600 m, although both methods have ra ther large uncer ta in t ies (Roedder and Bodnar, 1980). Colqui, and possibly Eureka, a r e thought t o have formed a t g rea t e r depths, approximately 1000 and 1400 m, respectively. Based only on geologic reconstruction es t imates , Round Mountain, Oatman, and DeLamar seem t o have formed a t shallower depths, possibly a s shallow a s 100 m. The shallow paleodepths and high gold/silver ra t ios of Round Mountain and O a t m a n a r e consistent with models for gold deposition presented by Henley (1985, this volume), Silberman and Berger (1985, this volume), and Hedenquist and Henley (1985).

CHAPTER 7

Sources of fluids--The predominance of me teo r i c wa te r has been documented in several epi thermal deposits, but only Creede (Bethke and Rye, 19791, Colaui (Kamilli and Ohmoto. 1977). and Sunnyside ( ~ a & d e v a l l and Ohmoto, 1977) habe had isotopic s tudies re la ted t o paragenet ic sequences. However, even with these more comprehensive studies, t h e sys t ema t i c s of hydrogen and oxygen isotopes may be more complex than previously thought. Very deta i led sampling of quar tz crys ta ls a t Creede by Foley e t al. (1982) has shown t h a t t h e previously reported 6 D values fo r quar tz (Bethke and Rye, 1979) do not t ru ly represent primary fluids, but ra ther represent a mixture, during extract ion and analysis, of hydrothermal fluids trapped in primary fluid inclusions, and isotopically l ighter, overlying ground waters, t rapped in pseudosecondary inclusions (discussed more thoroughly later). Without such detailed data , i t is difficult t o draw any more specific conclusions other than t h a t t h e mineralizing solutions were deeply circulating, dominantly me to r i c waters. I t should be noted, however, t h a t a magmat ic component of up t o ten percent could be hidden in t h e uncertainties and cannot be ruled out.

Source of sulfur and lead--For t h e 16 deposits considered by Heald e t al. (19861, t he sulfur-isotopic da ta fo r t h e Adularia-Sericite deposits a r e l imited t o measurements made on Creede (R. 0. Rye, U.S.G.S., personal communication, 1985), Sunnyside (Casadevall and Ohmoto, 19771, and Colqui (Kamilli and Ohrnoto, 1977). Although t h e 6 3 4 ~ values for t h e sulfides fo r a l l t h r e e deposits c lus ter very close t o 0 permil, qui te d i f ferent in terpre ta t ions of t he sources of t he sulfur have been proposed for t h e d i f ferent deposits. This is due i pa r t t o geologic considerations, but mainly t o 'I t he 6 4~ values for t h e sul fa te minerals; if equilibrium conditions a r e assumed, the re must be a mater ia l balance between the isotopically light sulfides and heavy sulfates. For Sunnyside and Colqui, t h e sul fa tes range f rom + I 0 t o +25 permil, but a t Creede they vary f rom +17 t o +45 permil. The significance of t h e unusually heavy sul fa te sulfur a t Creede will be discussed later.

A t Sunnyside, Casadevall and Ohmoto (1977) suggest t h a t t h e sulfur was derived f rom evaporite- bearing s e 'mentary rocks which have a bulk sulfur

$4 - value of 6 S - +12 permil a s sul fa te and a r e located outside of t he San Juan and Silverton calderas. They base thei r conclusion on t h e assumptions of sulfide- sul fa te isotopic equilib ium (based on a pyrite- anhydr i te pair having a 6 j4S di f ference of 22.3 permil) and t h a t t o t a l sul fa te concentration in the fluid was g rea te r t han or equal t o to t a l sulfide concentra t ion (indicated from mineralogical data). In order t o produce t h e 0 perrnil sulfides in equilibrium with t h e heavier sulfates a t t he presumed equilibration t empera tu re of 300°c, a source of 6 3 4 ~ = +12 permil is required. Kamilli and Ohrnoto (1977) also prefer a sedimentary sul fa te source for t h e sulfur a t Colqui, but suggest t h a t an igneous origin i s possible.

A t Creede, i t is c lear t h a t sulfate-sulfide isotope relationships were not governed by equilibrium fractionation, and t h a t aqueous sulfide and su l f a t e were essentially independent systems (Bethke e t al., 1973). Although t h e uniquely heavy sul fa tes a t Creede

require extensive isotopic evolution, Bethke et al.'s proposal for non-equilibrium, essentially independent sulfide-sulfate sys tems is also tenable fo r t h e o ther deposits. I t is probable t h a t non-equilibrium sulfide- sul fa te conditions exis ted for most epi thermal deposits (Ohmoto and Lasaga, 1982). Ra the r than assuming sulfide-sulfate equilibrium and relying on sca rce sedimentary units, sulfate/sulfide ratios, and mater ia l balance considerations t o produce 0 permil sulfides, i t is perhaps be t t e r t o assume t h a t t h e isotopic composition of the sulfide sulfur was controlled by the thick volcanic piles associated with e a c h of these deposits.

For those Adularia-Sericite d is t r ic ts where lead- isotopic studies have been done, Tonopah (Zartman, 1974), Pachuca (Cumming e t al., 1979) Creede, Lake Ci ty , Colorado, and Sunnyside (Doe e t al., 19791, t h e relatively radiogenic initial-lead ra t ios of galena suggest t h a t a significant component of t h e lead may have been derived f rom Precambrian o r Phanerozoic rocks underlying t h e volcanic rocks. This indicates t h a t t h e o re components of Adularia-Sericite deposits may have been derived, in large par t , by leaching of wallrocks deep in the system.

Character is t ics of Acid-Sulfate-type Deposits

St ructura l setting--Four of t h e five Acid-Sulfate deposits l isted in Table 7.2 a r e spatially r e l a t ed t o t h e margins of calderas. The other Acid-Sulfate deposit, Julcani, is associated with sil icic domes a t t h e in tersect ion of major faults. The association of these deposits with intrusive centers , particularly ring- f r ac tu re volcanic domes on t h e margins of calderas, appears t o be a cr i t ica l genet ic f ac to r (Heald e t al., 1986) in contras t t o the Adularia-Sericite-type deposit where t h e role of calderas is one of ground preparation for l a t e r hydrothermal fluids (Lipman e t al., 1976; Steven and Lipman, 1976).

Size of deposit--On t h e whole, Acid-Sulf a t e deposits a r e smaller in t e rms of tonnage than Adularia-Sericite-type deposits, although this may be a resul t of a l imited d a t a base. In addition, t h e surface projection of the productive a reas is relatively equidimensional ra ther than elongated. For example, t he a rea l ex ten t of t h e most productive par t of Goldfield, Nevada, is only approximately 2 km long by 1.5 km wide. For a l l t he Acid-Sulfate-type deposits, t h e ver t ica l e x t e n t of t h e mineralized a r e a i s much smaller than the horizontal ex ten t , usually less than 500 meters.

Host rock--The primary host rock for t h e Acid- Sulfate-type deposit is a lmost exclusively rhyodacite which is commonly porphyritic. A t Julcani, some o re also occurs in daci te , and a t Goldfield, t rachyandesi te and rhyolite, in addition t o rhyodacite, host t h e ore. The t iming of o re deposition re la t ive t o the emplacement of t he host rock is of particular importance. Age da te s show t h a t o r e deposition very closely ( 0.5 m.y.1 followed emplacement of t he host rock (Table 7.3) indicating a possible genet ic relationship. An exception is t he Lake C i ty I1 deposit (i.e., t h e Lake district; s ee foo tno te 5 on Table 7.2) where the ores occur in a quar tz l a t i t e ash-flow tuff which is considerably older than t h e ore. However, t he

D. 0. HAYBA, P. M. BETHKE, P. HEALD, & N. K. FOLEY

Table 7.3--Age and type of host rock and age of mineralization for five Acid-Sulfate districts (modified from Heald et al., 1986).

Age of Host Rock Age of District Host Rock (m-y. Mineralization Reference

Red Mtn., rhyodacitedome 21.3t023.6 23.1f0.6 Lipman et al., 1976 Colo. Mehnert et al., 1973

Gilzean et al., 1984

Julcani, rhyodacite dome 9.67 f 0.05 9.83 f 0.08 Petersen et al., 1977 Peru dacite dome to 10.13 f 0.03 Noble and Silberman,

1984

Lake City 11, quartz latite (28.4, >27.8 22.5l Slack, 1980 Colo. ash-f low Lipman et al., 1976

Summitville, rhyodacite dome 22.8 f 0.6 22.4 f 0.5 Mehnert et al., 1973 Colo.

Goldfield, rhyodacite dome 21.3 f 0.3 21.0 f 0.4 Silberman and Ashley, Nev . trachyandesite 21.5f 0.5 1970

rhyolite 28-33

'slack (1330) proposes that these later ores were generated during emplacement of the nearby Red Mountain rhyodacite dome, approximately 22.5 m.y. ago.

ore a t Lake C i ty I1 fills f r ac tu res t h a t form a radial pa t t e rn around a rhyodacite flow dome t h a t Slack (1980) suggested is genetically re la ted t o t h e Lake C i ty I1 ores and has a similar age t o t h e ores. Hon e t al. (1985) have shown t h a t some veins included in t h e Lake C i ty I1 district a r e older than t h e rhyodacite flow dome and probably formed from an ear l ier hydrothermal system(s).

Mineralogy--Acid-Sulfate-type deposits a r e character ized by t h e occurrence of t h e vein mineral assemblage enargi te + pyrite +/- covellite. Adularia and chlor i te a r e absent or rarely present. Ore occurs primarily a s native gold and e lec t rum with sulfides, sulfosalts, and tellurides. Bismuthinite has been identified in 3 of t he 5 Acid-Sulfate deposits in Table 7.2, but selenides, rhodochrosite, and f luor i te a r e rare.

Metal ratios--Summitville and Goldfield have low silver-to-gold ratios (2 : l ) which ref lec t t h e high proportion of f r e e gold and gold-bearing minerals. Julcani, Red Mountain and Lake C i ty I1 have high silver-to-gold ra t ios ( 10:l) and a r e cha rac te r i zed by more abundant silver mineralization, primarily in t h e form of silver sulfides and sulfosalts. Copper const i tu tes a major proportion of the base-metal production for Acid- Sulfate-type deposits, especially a t Goldfield and Summitville, where copper accounts for more than 85% of the base-metal production. The silver-rich dis t r ic ts have proportionally m o r e base metals, but copper production is secondary t o lead (1:2 Cu:Pb ratio).

Wallrock alteration--A definit ive cha rac te r i s t i c of Acid-Sulfate-type deposits is t h e association of advanced-argillic a l tera t ion with the ore. Kaolinite, usually accompanied by alunite, occurs close t o t h e

vein and is of ten coextensive with silicification. Fa r the r f rom t h e vein, argil l ic a l tera t ion, somet imes intermixed with ser ic i t ic a l tera t ion, surrounds the zone of advanced-argillic a l tera t ion. The argil l ic a l tera t ion zone is commonly mineralogically zoned, with kaolinite nearer t h e vein and smec t i t e f a r the r from the vein. The outermost a l t e ra t ion zone consists of propylitic alteration.

Thermal history--The l imi ted fluid-inclusion d a t a for Acid-Sulfate-type deposits indicate t h a t o re deposition occurred primarily a t t empera tu res similar t o those of Adularia-Sericite-type deposits ( 2 0 0 ' ~ t o 3 0 0 ~ ~ ) . The salinit ies of Acid-Sulfate-type deposits, which a r e not yet well documented, show a wide range. Secondary inclusions in qua r t z phenocrysts which a r e thought t o conta in t h e fluids responsible for t he in tense quar tz + aluni te +/- pyr i te a l tera t ion which preceded o r e deposition have salinity ranges of 5-24, 7-21 and 5-18 wt.-% NaCl equivalent a t Julcani, Summitvil le and Goldfield, respectively (Bruha and Noble, 1983). However, l imi ted d a t a from primary inclusions in quar tz and sphaler i te associated with main s t age o re deposition f rom Lake C i ty I1 (Slack, 1980), Red Mountain (Nash, 19751, and Summitvil le (Stoffregen, 1985) indicate salinit ies on the order of 1 t o 6 wt.-% NaCl equivalent. These l imited d a t a a r e examined more closely in a la ter section, but clearly, much remains t o be learned f rom fluid-inclusion studies of Acid-Sulfate-type deposits.

Paleodepth--Paleodepth e s t ima tes for Acid- Sulfate-type deposits appear t o be similar t o Adularia- Ser ic i te deposits (300-600 m), although Nash (1975) suggests t h a t t h e Red Mountain deposit may have formed a t depths g rea te r than 1200 m. Based solely on

geologic reconstruction, t he gold-rich deposits, Goldfield and Summitvil le, may have formed a t somewhat shallower depths than the silver-rich deposits.

Source of constituents--The stable-isotope ciagt a on Acid-Sulfate deposits i s l imited t o Xqme 6 0 numbers fo r Goldfield (Taylor, 1973) and 6"s da ta for Julcani (Goodell. 1970). Goldfield (Jenson e t al.. 1971). and ~ u k m i t v i l i e (R.'o. Rye, U.S.G.S., personai communication, 1985). Without any deuterium data , t h e source of t he hydrothermal fluid is equivocal. The possibility of a significant magmat ic component is discussed later. The pauci ty of sulfur-isotope d a t a permits only generali t ies. The sulfides from these th ree districts plot between -7 and + 3 permil suggesting a magmat i c source for t he sulfur. Similarly, lead-isotopic s tudies a t Summitville and Red Mountain show t h a t t h e galena is very similar isotopically t o the enclosing volcanic rock, implying t h a t e i the r t h e adjacent rocks or related magmat ic fluids were the primary source of the lead (Doe e t al., 1979).

Summary of Character is t ics -- The primary character is t ics which distinguish the

Adularia-Sericite-type deposits f rom t h e Acid-Sulfate- type deposits a r e t h e vein mineralogy and the a l tera t ion assemblages (Heald e t al., 1986). As shown in Table 7.1, t he re a r e also many other important, but less definit ive, character is t ics of each type which need to be considered in developing genet ic models. A comparat ive study, such a s t h e one done by Heald e t al. (1986) helps de te rmine which character is t ics a r e t h e sa l ient f ea tu res of a deposit-type model and which f ea tu res a r e just local variations. In order t o examine e a c h of these two types of epi thermal deposits in greater detail , a character is t ic deposit from each group will be reviewed in t h e next sections.

THE ADULARIA-SERICITE ENVIRONMENT: CREEDE AS AN EXAMPLE

In t h e foregoing sect ion t h e character is t ics of 16 well-documented Ter t iary volcanic-hosted epi thermal ore deposits were summarized based on the comparat ive study done by Heald e t al. (1986). I t was concluded t h a t t w o distinctive types of deposits could be distinguished: I ) those deposits character ized by a n a l tera t ion assemblage dominated by adularia and ser ic i te , and 2) those deposits character ized by an a l tera t ion assemblage containing kaolinite and alunite. The most thoroughly studied of t h e Adularia- Sericite-type deposits i s t he Creede mining district , Colorado. Therefore, Creede is useful a s t h e exemplar for t he Adularia-Sericite-type deposits.

Creede a s a n Exemplar

In basing our discussion of the Adularia-Sericite environment on Creede we recognize t h a t some fea tu res of t he Creede district a r e not representa t ive of t h e group a s a whole. For example, t h e chemical and isotopic evolution of the Creede o re fluids in a

closed-basin lake prior t o incorporation in to the hydrothermal system is an aspect t h a t may be uncommon, although a similar geologic se t t ing may have exis ted a t Calico, California. The presence of an evaporat ive lake is obviously not a requirement for o re format ion in these deposits, although the high salinit ies developed by evaporation probably made the Creede sys tem highly ef f ic ient in the transport of base-metals (cf. Henley, 1985, this volume). We also recognize t h a t within this group of deposits t he re is considerable diversity in specific characterist ics. For example, Round Mountain and Oatman a r e both distinctively gold-rich and sulfide-poor re la t ive t o t h e o ther d is t r ic ts discussed in t h e previous section. Other

I 1 COLORADO ( I

Known or readily Calderas w i t h major inferred calderas associated mineral deposits

Figure 7.1. Calderas of the San Juan volcanic field: S, Silverton; LC, Lake City; CP, kchetopa Park; Bz, Bonanza; U;, La Garita; SL, San Luis; B, Bachelor; C, Creede; MH, Mount Hope; PI Platoro; Stl, Summitville; L, Lost Lake; U, Ute Creek; SJ, San Juan; UN, Uncompahgre; MI general location of the Elammoth Mountain caldera. Af ter Steven and Eaton (1975) .

D. 0. HAYBA, P. M. BETHKE, P. HEALD, & N. K. FOLEY 137

Figure 7.2. Generalized geology of the Creede and San Luis calderas in relation to remnants of the Bachelor (B) and La Garita calderas and to the Creede mining district (shown in box). Control is moderate to good where boundaries are shown by solid symbols, and conjectural where shown by open symbols. Pr, Point of rocks volcano; S, Spar City. After Steven and Lipman (1976).

dif ferences could be enumerated. In spi te of t h e diversity in speci f ic characterist ics, we would argue t h a t t h e d i f ferences between the various d is t r ic ts represent variations on a theme, whose main character is t ics a r e i l lustrated by the similarities.

We a r e aware t h a t t he "Hot-Spring-type" deposits such a s McLaughlin, California, and Hasbrouck Mountain and Sulphur, Nevada, a r e not specifically t r ea t ed when Creede is used a s a n exemplar because t h e r e is no evidence of surface discharge of t h e o r e fluids a t Creede. I t is our current opinion, however, t h a t such hot-spring-type deposits form in t h e surficial pa r t s of hydrothermal sys tems similar t o those which deposited t h e Creede ores. Berger and Eimon (1983), Henley (1985, this volume), Silberman and Berger (1985, this volume), and Berger and Silberman (1985, this volume) t r e a t hot-spring- type deposits separa te ly because of important s t ructura l a t t r ibu te s tha t , in t e rms of o r e controls, somewhat distinguish these deposits f rom t h e deeper classical vein deposits. However, Silberman and Berger (1985, this volume) also present evidence f rom Bodie, California, t h a t links hot-spring-type deposits directly t o t h e Creede-type veins.

Summary of Important Studies

The Creede mining dis t r ic t has been t h e focus of both extensive and intensive study fo r the past 30 years. Studies by T. A. Steven and co-workers of t h e U.S.Geologica1 Survey unravelled t h e volcanic history of t h e district , re la ted t h e o r e deposits t o t h a t history, and documented t h e s t ructura l control of o re

deposition (Steven and Ra t t6 , 1965; Rat tC and Steven, 1967; Steven and ~ a t t 6 , 1973; Steven and Liprnan, 1973; Steven, 1967; Steven and Eaton, 1975). Continued studies by T. A. Steven and P. W. Lipman and co-workers have worked out t h e volcanic s t ra t igraphy and volcano-tectonic evolution of the en t i r e San Juan volcanic field, providing a particularly well-developed regional context for t he Creede dis t r ic t (Steven and Lipman, 1976; Lipman e t al., 1970; Lipman et al., 1978; and Doe et al., 1979). Detailed mineralogical-geochemical studies have served t o develop a well-documented, comprehensive model of ore deposition (Roedder, 1960; Roedder, 1965; Roedder, 1977; Bethke e t al., 1976; Barton e t al., 1977; Wetlaufer, 1977; Bethke and Rye, 1979; Woods e t al., 1982; Foley e t al., 1982; Heald-Wetlaufer and Plumlee, 1984; Hayba, 1984; Robinson and Norman, 1984; Plumlee and Hayba, 1985). These studies have been complemented, and the models developed f rom them improved, by detailed, unpublished geologic studies by t h e Homestake Mining Company, Pioneer Nuclear Corporation, Chevron Resources, Freepor t Mining Company, Minerals Engineering Company, and by a number of theses (Cannaday, 1955; Chaffee , 1967; Hull, 1970; Giudice, 1980; Robinson, 1981; Battory, 1982; McCrink, 1982; Wason, 1983; Rice, 1984; Horton, 1983; Vergo, 1984; and Misantoni, 1985).

Geologic and Mineralogic Character is t ics

Volcanic-hydrothermal history--The San Juan Mountains a r e the main erosional remnant of a volcanic field which covered most of t h e southern

Rocky Mountains in mid-Tertiary t ime. Approximately 213 of t h e volume of the present field is composed of a ser ies of ear ly lavas and volcaniclastic aprons, mainly of andesi t ic composition, re la ted t o a number of s t ra tovolcanic centers. These "Early Intermediate" composition lavas a r e overlain by a ser ies of qua r t z l a t i t i c t o rhyolit ic ash-flow sheets. F i f t een calderas have been identified a s sources of these ash-flows (Fig. 7.1). Many of these ca lderas a r e nes ted to form complexes such a s the cen t r a l San Juan caldera complex which hosts t h e Creede mining district .

The Creede mining dis t r ic t is located along the western edge of t h e centra l San Juan caldera complex, a ser ies of 7 nes ted calderas from which quar tz- la t i t ic t o rhyolit ic ash-flow shee t s erupted over t h e brief in terval of 27.6 to 26.2 m.y. (Fig. 7.2). Only 5 of t he 7 calderas can b e accurate ly located, t h e remaining t w o having caved in to the younger Creede caldera. The Creede o res a r e hosted by t h e in t racaldera fill of t he resurgent Bachelor caldera, t he second caldera in the series. The o res a r e contained in a s e t of f r ac tu res comprising a graben running between t h e Creede caldera, t h e youngest of t h e series, and t h e slightly older San Luis caldera.

Radiometr ic dating of vein adularia and mixed- layer clay a l tera t ion minerals (Bethke et al., 1976; Vergo, 1984) indicates t h a t t h e ores were deposited approximately 1 million years a f t e r t h e youngest dated volcanic event in the district . These da te s a r e consistent with t h e observation of Steven and R a t t e (1960b) t h a t mineralization was confined t o s t ruc tu res young enough t o c u t t h e sediments of t h e Creede format ion which fi l l t he moa t of t he resurgent Creede caldera.

Studies of the Creede Format ion by Steven and R a t & (19651, Steven and Van Lonen (19701, McCrink (1982), and Battory (1982) have shown t h a t t h e sediments filling the moa t of t h e Creede caldera consist of landslide debris and s t r eam channel fi l l on the margins of t h e moat, and lacustrine deposits of airfall and water-lain tuffs, interbedded with a f ew thin ash-flow tuffs, in t h e cen te r of t h e moat. These sediments accumulated in a shallow, playa lake environment, and the lacustrine sediments were strongly zeolit ized during diagenesis.

Steven and Eaton (1975) proposed t h a t t he convecting hydrothermal sys tem was driven by an unexposed pluton beneath t h e Creede district . They observed t h a t t he maximum displacement along the Amethyst and Bulldog faul t sys tems occurs in t h e central , most heavily mineralized pa r t of t h e district , where subtle magnet ic and gravity anomal ies a r e suggestive of a buried intrusion.

Creede o re deposits--The base- and precious- me ta l mineralization a t Creede fills open f rac tures . The o re zone, which occupies a narrow ver t ica l range from 250 t o 400 m, has been mined nearly continuously for approximately 3 km along s t r ike on t h e Amethyst- OH vein system, and for over 2 km on t h e Bulldog Mountain vein system (Fig. 7.3). This horizontally- dominated aspect ra t io is character is t ic of deposits of t h e Adularia-Sericite-type. Mineralization is confined to the f r ac tu re system comprising the Creede graben bounded on t h e e a s t by the Solomon-Holy Moses faul t system and on t h e west by t h e Alpha-Corsair system

(Fig. 7.3). This graben sys tem follows closely t h e s t ructures of an older graben in terpre ted t o be the keystone graben of t h e resurgent Bachelor caldera. Detailed o re petrology studies of mater ia l f rom t h e major producing s t ruc tu res (Amethyst, OH, P and Bulldog Mountain vein sys tems) indicate t h a t a l l were par t of a single hydrologic system and were filled during t h e same mineralizing even t (Barton e t al., 1977; Bethke and Rye, 1979; Heald-Wetlaufer and Plumlee, 1984).

A significant, low-grade, bulk-tonnage silver resource also exists a s disseminated replacements in the wallrocks of t he upper par ts of the Amethyst vein system near i t s in tersect ion with t h e OH vein (Giudice, 1980). Additional bulk-tonnage resources a r e present a s mineralized s t r eam sediments in t h e s t r eam channel fac ies of t he Creede Format ion adjacent t o i t s

Figure 7.3. Generalized geology of the Creede mining district, modified from Steven and Eaton (1975) . Area of map is shown in Figure 7.2. Faults are dashed where uncertain or inferred; bar and ball show the downdropped side. Tc, Creede Formation; Tf, Fisher quartz latite flow; Tfi, volcanic neck of Fisher quartz latite.


NORTH SOUTH

present surface

? OH section Amethyst section

OH assemblege

F i g u r e 7.4. G e n e r a l i z e d l o n g i t u d i n a l p r o j e c t i o n o n t o a v e r t i c a l p l a n e o f t h e OH v e i n a n d the s o u t h e r n end o f the Amethyst vein. The main workings a n d the g e n e r a l d i s t r i b u t i o n o f the OH a n d B u l l d o g a s s e m b l a g e s are shown. The " s t r u c t u r a l d i s c o n t i n u i t y " marks the p o s i t i o n t h a t would have b e e n the i n t e r s e c t i o n o f the OH and Amethyst v e i n s had f r a c t u r i n g b e e n more c o n t i n - uous . T h e A m e t h y s t v e i n has b e e n m i n e d n o r t h w a r d f o r a b o u t 2 km b e y o n d the " s t r u c t u r a l d i s c o n t i n u i t y " but is n o t shown here. Adapted f rom B a r t o n et al. (1977).

truncation by the Amethyst vein (Wason, 1983; Rice, 1984). The meta ls have been introduced into the c las t ic fac ies of t h e Creede F o r ~ n a t i o n by leakage of the o r e fluid from underlying or adjacent veins.

Mineralogic characterist ics--Mineralization in the Creede veins is strongly zoned from an association in t h e north (OH Assemblage) of chlor i te + hemat i t e + quar tz + adularia + sphaler i te + galena + chalcopyrite + pyrite +I- fluorite and t e t r ahedr i t e t o a bar i te + rhodochrosite + quar t z + adularia + galena + sphalerite + f luor i te + t e t r ahedr i t e + silver sulfosalt + nat ive silver association t o t h e south (Bulldog Assemblage) (Fig. 7.4). Heald-Wetlaufer and Plumlee (1984) have dernonstrated t h a t t h e Bulldog Mountain and OH-P- Amethyst vein sys tems e a c h contain both the OH and Bulldog assemblages, and t h a t t he two a r e contemporaneous and re la ted t o e a c h other through facies changes along t h e fluid-flow path. The mineralization in each of these f ac i e s may be divided in to 5 s tages on t h e basis of mineralogy and texture , and e a c h of the s tages can be corre la ted between the t w o assemblages. Deta i ls of t h e mineralogy and t h e correlation of s tages between t h e t w o assemblages a r e given in Table 7.4.

A l a t e association of covelli te + chalcoci te + acan th i t e has been described by Giudice (1980) and Robinson and Norman (1984) in t h e disseminated ore from t h e upper and southern Amethyst vein system. This assemblage is a volumetrically minor, but economically important , component of t h e Bulldog Assemblage. I t has not yet been incorporated in to t h e geochemical model for Creede ore deposition, because i t s l imited occurrence and fine-grained nature makes i t difficult t o study.

Many minerals of t he Creede o res (sphalerite, rhodochrosite, siderite, Fe-chlorite, te t rahedr i te- tennant i te , proustite-pyrargyrite, l a t e gel-pyrite, and, probably, o ther sulfosalt minerals) exhibit marked compositional variations in both t ime and space. Of these minerals, only sphalerite (Barton, e t al., 19771, t he Mn-Fe-carbonates (Wetlaufer, 1977) and Fe-

chlor i te (Horton, 1983) have been studied extensively. The co~nposi t ional variations in the Mn-Fe carbonates and Fe-chlorite have proven too complex t o be very useful t o developing a gene t i c model of Creede, but those in t h e sphaler i te (primarily iron content) have provided t h e main basis for documenting the evolution of t h e o re fluid in t i m e and space, and for constraining the chemical environment. The sphalerite is beautifully banded, reflecting, mainly, t h e variations in iron content. This banding records a wealth of information about crys ta l growth and dissolution. Even more importantly, t he growth zones can be corre la ted along the OH vein and between the OH and other veins. Figure 7.5 shows a composite microprobe t rac ing of samples of sphaler i te f rom t h e OH vein t h a t is representa t ive of most of t he B and D s t age paragenesis.

Hydrothermal leaching (intense for barite, f luorite and rhodochrosite, and less intense for sphalerite and galena) demonstra tes t h a t a t t imes t h e solutions enter ing t h e o re zone were undersaturated with respect t o se l ec t ed components. The increased intensity of leaching and number of leaching horizons a t t he north end of t h e OH vein suggest t h a t t he solutions en te red t h e o re zone in t h e north and t raversed t h e vein system from north t o south.

Supergene oxidation has a f f ec t ed t h e Creede ores to a l imited ex ten t producing extensive manganese oxide (presumably a f t e r rhodochrosite) in the southern Amethyst system and local occurrences of kaolinite and halloysite. Some of the oxidized rocks contain distinctive f r ac tu re coatings of dendrit ic nat ive silver. Veinlets of fine-grained alunite which crosscut mineralization in the upper levels of t he Amethyst vein system have been dated by potassium- argon techniques and range in age from 3.5 t o 5.0 million years (M. Lanphere, U.S.G.S., personal communication, 1981). This age corresponds closely t o the regional uplift of t h e southern Rocky Mountains which, presumably, raised the Creede orebodies above t h e water table. I t is no t c lear whether any or a l l of

T a b l e 7.4--Mineralogy and p a r a g e n e t i c s t a g e s of t h e OH and Bulldog assemhlages of t h e Creede d i s t r i c t based on t h e work of P. B. Barton and P. Y. Bethke a s r e p o r t e d i n Bethke and Rye (1979) f o r t h e OH, P, and Amethyst v e i n systems, on t h a t of Robinson (1981) and Giudice (1980) f o r t h e upper and s o u t h e r n Amethyst v e i n , and on t h e d e t a i l e d s t u d y by G. S. Plumlee and P. Heald on t h e Bulldog Mountain v e i n system a s r e p o r t e d i n Heald-Wetlaufer and Plumlee (1984).

S t a g e OH Assemblage Bulldog Assemblage S tage (OH, P n o r t h e r n Amethyst ( s o u t h e r n 2 / 3 Bulldog v e i n v e i n s and n o r t h e r n 1/3 system, southernmost OH v e i n Bul ldog v e i n sys tem) and sou the rn Amethyst v e i n )

E F ib rous p y r i t e wi th some F ib rous p y r i t e wi th minor V (younges t ) m a r c a s i t e and s t i b n i t e m a r c a s i t e , s t i h n i t e and s u l f o -

s a l t s . Late Ag (may b e supergene )

D R e l a t i v e l y coarse-grained Coarse-grained s p h a l e r i t e and IV s p h a l e r i t e , g a l e n a , chalco- ga lena wi th minor l a t e chalco- p y r i t e , and q u a r t z , some hema- p y r i t e , t e t r a h e d r i t e and Ag-Cu t i t e ; s i l v e r m i n e r a l s n o t a b l y s u l f i d e s and s u l f o s a l t s a b s e n t ; subdivided i n t o t h r e e s u b s t a g e s on b a s i s of c o l o r banding i n s p h a l e r i t e : i n n e r l i g h t , middle da rk , and o u t e r l i g h t

C Vo lumet r i ca l ly minor s i d e r i t e - Vo lumet r i ca l ly minor (Mn,Fe)- 111 manganos ide r i t e and f l u o r i t e ca rbona te and f l u o r i t e on whi t e on q u a r t z ; s i ts on deep e t c h and amethys t ine q u a r t z ; most of e a r l i e r B s t a g e ; most f l u o r i t e e t c h e d - commonly f l u o r i t e deep ly e t ched - comple te ly removed commonly complete ly removed

B R e l a t i v e l y f ine -g ra ined sphal- In te rbanded b a r i t e and f i n e - I I e r i t e , g a l e n a , c h a l c o p y r i t e , g r a i n e d s p h a l e r i t e , g a l e n a , and c h l o r i t e , h e m a t i t e , p y r i t e , t e t r a h e d r i t e and some t e t r a h e d r i t e - t e n n a n t i t e

A P r i m a r i l y q u a r t z w i t h minor E a r l y massive r h o d o c h r o s i t e , I ( o l d e s t ) c h l o r i t e and s u l f i d e l a t e r b a r i t e ; minor d i s semina ted

s p h a l e r i t e and g a l e n a

the wire and leaf native silver, which is cha rac te r i s t i c of t h e Bulldog assemblage, is of supergene origin. Our current in terpre ta t ion is t h a t i t is primary.

wallrock a l t e ra t ion - -~a l l rbck a l tera t ion a t Creede consists of a feldspar-destructive, mixed-layer i l l i te /smect i te a l tera t ion present along t h e top of t h e orebodies." Alteration intensity ranges f rom weak, where only t h e pumice f ragments a r e a l tered, t o intense, where the ent i re rock has been a l t e red and where all evidence of primary volcanic t ex tu re has been obl i tera ted (Horton, 1983). The a l tera t ion is strongly fracture-controlled, and the intensity of argil l ization decreases away from the veins.

*Heald e t al. (1986) included such mixed-layer illite- s m e c t i t e a l tera t ion in t h e ser ic i te ca tegory, and we will use the t e rm "sericite" in thei r sense to include both i l l i te and i l l i te-smectite mixed-layer clay minerals. -

The most in tense a l tera t ion forms a "clay cap" t h a t marks t h e upper l imit of mining. The ver t ica l e x t e n t of the clay c a p has not been well established, but surface outcrop above highly a l t e red a reas o f t en shows no a l tera t ion effects , and drilling above the OH

D. 0. HAYBA, P. M. BETHKE, P. HEALD, & N. K . FOLEY

] 1 SECTION CONTINUES TO THE RIGHT

-I - -- Oldest

70 1 Youngest

I SECTION CONTINUED FROM THE LEFT

MOLE PERCENT FeS in SPHALERITE

Figure 7.5. Coarse-scale plot of the iron content across a growth-banded aggregate of sphalerite crystals, modified from Barton et al. (1977). Data are taken from composite microprobe tracings of samples from the north end of the OH vein and are replotted to eliminate instrumental scatter. A). tracing of the iron content across finely banded, early (B-stage), undif- ferentiated sphalerite: B). tracing of the iron content across coarse, late (D-stage) sphalerite that is subdivided on the basis of color and iron and cadmium contents.

vein is repor ted t o have encountered only f resh rock approximately 100 m above the base of alteration. Along t h e OH vein, t h e clay cap is continuous and regularly distributed (Fig. 7.6). In o ther par ts of t h e district , t h e distribution of t h e clay cap is much more irregular; t he base of a l tera t ion undulates over a ver t ica l in terval of several hundred f ee t , and pa r t s of t he veins show no a l tera t ion a t a l l a t any exposed level. The continuity and regularity of t h e base of t h e clay c a p above t h e OH vein is in terpre ted t o r e f l ec t a more uniform hydrologic regime in the vertical , simple tension f r ac tu re t h a t contains t h e vein, in contras t t o t h e hydrology in t h e more complex s t ruc tu res of t h e o ther vein sys tems in the district . Barton e t al. (1977) in terpre ted t h e clay or ser ic i t ic a l tera t ion t o mark a

region of recondensation of acid-rich s t eam derived by boiling of t h e o re fluid.

In the upper par ts of t h e vein systems, particularly near t h e junction of t h e Amethyst and OH veins, t he wallrock adjacent t o t h e veins i s highly silicified and somet imes bleached. The sil icification appears t o pre-date t h e ser ic i t ic a l tera t ion and t o have protected t h e a f f ec t ed wallrocks from it. In a reas of in tense silicification, t he ser ic i t ic a l tera t ion o f t en borders t h e silicified zone, although t h e silicification may pass in to fresh, unsilicified wallrock, particularly in a reas lying beneath t h e ser ic i t ic cap. I t is suggested, but not demonstrated, t h a t t h e sil icification predated t h e ser ic i t ic a l tera t ion and protected t h e walls from t h e a l ter ing fluids.

142 CHAPTER 7

EXPLANATION

lllite alteration Boiling, fluid inclusions

Present surface * Adularia localities

- OH section - - Amethyst section-

NORTH 0 200 400 METERS - SOUTH

0 800 1600 FEET

F i g u r e 7.6. G e n e r a l i z e d l o n g i t u d i n a l p r o j e c t i o n o n t o a v e r t i c a l p l a n e o f the OH v e i n a n d the s o u t h e r n end o f the Amethyst vein. The base o f the i n t e n s e sericitic a l t e r a t i o n is i n d i c a t e d : t h e p o s i t i o n o f the top is unknown.

Detailed s tudies of t h e variation in s t ructure and composition of t h e mixed-layer clay minerals by Horton (1983) and Vergo (1984) show tha t t h e i l l i te component of t h e clays ranges f rom greater than 95% t o less than 60%. In t h e a r e a just north of t h e Amethyst-OH intersection, where the ground between t h e t w o is c u t by many smal ler veins, t he proportion of i l l i te in the mixed-layer clays increases toward t h e Amethyst vein. The degree of ordering of t h e stacking sequence of t h e i l l i te and s m e c t i t e layers, which ranges from random t o long-range ordered (Reichweite = 0 t o 31, also increases toward t h e Amethyst vein. Using empirical relationships between t empera tu re and t h e s t ructure and composition of mixed-layer illite- smec t i t e clays, Horton (1983) e s t ima ted the position of isotherms surrounding t h e upper portions of t h e Amethyst vein near i t s in tersect ion with the OH vein. Vergo (1984) found t h e mixed-layer clays along the Bulldog vein system t o be more smectite-rich than most of Horton's samples (presumably because they formed a t lower temperatures). Vergo did not find any sys temat ic variation in t h e i l l i te-smectite r a t io with distance from t h e vein for d is tances up t o 70 m from the vein, and calcula ted t h a t t h e system must have been act ive for a period of a t l ea s t 10,000 years t o produce such an apparent ly f l a t t he rma l gradient.

Below the zone of in tense ser ic i t ic a l tera t ion, t h e wallrocks enclosing t h e veins a r e essentially fresh, but were enriched in potassium by a period of potassium feldspar-stable, metal-barren hydrothermal a l tera t ion t h a t occurred approximately 2 million years prior t o ore deposition ( R a t t i and Steven, 1967; Bethke e t al., 1985). This in tense potassium metasomat ism was presumably re la ted t o hydrothermal circulation in the keystone graben of the Bachelor caldera, and is similar in most respects t o t h e potassic a l tera t ion a t Bodie (OINeil e t al., 1973; Silberman and Berger, 1985, this volume). I t does not appear t o have been associated with any ear l ier period of mineralization.

geochemical environment of t h e Creede ore-forming system has been developed through fluid-inclusion, light-stable isotope, and lead-isotope studies, and through thermochemical analysis of t he o re and gangue mineral assemblages. A summary of t h e general character is t ics of t h e geochemical environment a r e shown in Table 7.5.

Temperature-salinity ranges--Much of the evidence used in defining t h e depositional pa ramete r s of t h e Creede sys tem has c o m e f rom the study of fluid inclusions. Numerous homogenization measurements have defined a t empera tu re range of 120' t o 2 8 0 ' ~ for t he Creede ore-forming fluids. Most of t h e measurements have been on sphalerite, qua r t z and fluorite f rom t h e OH-Amethyst vein system (Woods e t al., 1982; Robinson and Norman, 1984; J. Goss, U.S.G.S., personal communication, 1985) and on bar i te from the mineralized s t r eam channel of t h e Creede Formation near i t s t runcat ion by the Amethyst vein (Rice, 1984). There is a c lear t rend from higher t empera tu res in t h e nor th (OH data , a s summarized by Woods et al., 1982) t o lower t empera tu res in t h e south (southern Amethyst da t a , Robinson and Norman, 1984, and Creede Format ion da ta , Rice, 1984) although t h e t empera tu re ranges fo r minerals and veins overlap (Fig. 7.7).

Fluid-inclusion s tudies have also shown t h a t t h e Creede ores were deposited from relatively concentra ted NaCl brines and t h a t these brines mixed with overlying ground wa te r in t h e o r e zone (discussed below). From freezing and crushing and leaching studies, Roedder (1963, 1965) showed t h a t t h e o re fluid ranged in salinity f rom about 4 t o 12 wt.-% NaCl equivalent and had ave rage a tomic ra t ios of Na:K:Ca of 9:1:2. These ra t ios a r e in excel lent agreement with those ca lcula ted fo r a t empera tu re of 2 6 0 ' ~ using t h e alkali geothermometer of Fournier and Truesdell (1973). As noted in t h e previous section, with t h e exception of Colqui, Peru, the salinity of t he Creede ore fluids is much hinher than all t h e other Adularia-

Geochemical Environment Sericite-ty pe deposits-evaluated by Heald e t al. (1 986). Depth-pressure ranges--In addition t o the

An extremely detailed knowledge of the t empera tu re and salinity da ta , fluid-inclusion studies


Tab le 7.5--General env i ronmen ta l pa rame te r s f o r t h e OH v e i n , Creede , Colorado (mod i f i ed from Bar ton e t a l . , 1977)

Range Refe rence Source of Pa rame te r observed environment i n f o r m a t i o n

Temperature 190 - 2 8 5 ' ~ 2 5 0 ' ~ 40 - 50 b a r s 50 b a r s

Depth 450 - 600 m 500 m

S a l i n i t y 4 - 12 wt% 6 w t % Na:K 7.4 - 9.9 9

T o t a l S 0.018-0.30 mola l 0.02 mola12

F l u i d i n c l u s i o n s 1

Evidence of b o i l i n g i n f l u i d i n c l u s i o n s

Es t ima ted from p r e s s u r e and g e o l o g i c r e c o n s t r .

F l u i d i n c l u s i o n s Analyses of f l u i d

i n c l u s i o n s Analyses o f f l u i d

i n c l u s i o n s C a l c u l a t e d

' ~ o s t o f t h e f l u i d i n c l u s i o n ev idence i s from t h e l a t e r h a l f of t h e m i n e r a l i z a t i o n which i s much c o a r s e r g ra ined .

2 ~ e c a u s e o f t h e problem of s u l f u r c o n t r i b u t e d by o x i d a t i o n o f s u l f i d e s d u r i n g sample h a n d l i n g , t h e lower t o t a l s u l f u r v a l u e s a r e c o n s i d e r e d more r e l i a b l e .

have shown t h a t t h e o re fluid was, a t t imes , boiling near t h e top of t h e OH vein ( the evidence for boiling is discussed later) . Using t h e tables f rom Haas (1971), t h e pressure a t t h e top of t h e o re body was approximate ly 40 bars, based on t h e boiling point of a 250°c, 1 molal NaCl fluid. For mos t of o r e deposition, t h e pressure must have been slightly g rea t e r than this, because boiling appears t o have occurred infrequently. A hydros ta t ic head of approximately 450 m is necessary t o produce t h e requisite pressure, which is in excel lent ag reemen t with t h e geologic reconst ruct ion by Steven and Eaton (1975).

Chemical parameters--In 1977, Barton e t al. did a thermochemical analysis of t h e Creede sys tem and were able t o put l imi ts on such pa rame te r s a s ac t iv i t ies of S2 and 0 2 , pH, and to t a l sulfur. In t e r m s of t h e activities of S2 and O2 (Fig. 7.81, t h e Creede environment is located nea r t h e junctures of t h e hemat i te , pyr i te and Fe-chlorite stabil i ty f ields based on t h e common occu r rence of t h a t mineral assemblage and on t h e iron con ten t of sphaler i te which usually ranged between 1 and 4 mole-% FeS (Fig. 7.5).

Because of t h e common occurrence of adular ia associated with minor amoun t s of s e r i c i t e in t h e wall rock near t h e vein, Barton e t al. (1977) e s t ima ted t h e pH of t h e fluid during mineral deposition in t h e Creede sys tem based on t h e feldspar hydrolysis reac t ion

A pH of 5.4 is e s t ima ted f rom t h e thermodynamic d a t a of Montoya and Hemley (1975) fo r a 1 molal solution with a Na/K ra t io of 9. This i s a nearly neut ra l pH a t 2 5 0 ' ~ .

Roedder e t al. (1963) measured to t a l sulfur concentra t ions of about 0.02 molal in inclusion fluids

QUARTZ & BARITE

Figure 7.7. Histograms of homogenization temperatures for fluid inclusions in quartz and sphalerite from the OH vein (Woods et al., 1982), quartz from the southern Ame- thyst vein (Robinson and Norman, 1984), and barite from the Creede Formation (Rice, 1984).

F i g u r e 7.8. Log aS2-%2 d iag ram a t 2 5 0 ~ ~ show- i n g the m i n e r a l s t a b i l i t y f i e l d s f o r the s i g n i f i c a n t m i n e r a l s i n the C r e e d e ores. The shaded f i e l d o f m a g n e t i t e is c o m p l e t e l y p reempted by chlorite. The c o n t o u r f o r 20 m o l e p e r c e n t F e S i n sphalerite c o i n c i d e s w i t h the pyrrhotite f i e l d boundary. Quartz is p r e s e n t t h r o u g h c u t the diagram. Abbre- v i a t i o n s : py = pyrite, ccp = chalcopyrite, b n = b o r n i t e , h e m =hematite, chl = chlo- ri te. The s t a n d a r d s t a t e f o r S 2 a n d O2 is the i d e a l d i a t o m i c g a s a t 1 a tmosphere and 2 5 0 ~ ~ . The data f o r the i r o n - c o p p e r s u l - f i d e s a n d o x i d e s are s u m m a r i z e d i n B a r t o n a n d S k i n n e r ( T a b l e 7.2, 1 9 7 9 ) ; t h i c k sol id l i n e s : b o u n d a r i e s b e t w e e n p y r i t e , p y r r h o - t i t e , chlorite, a n d hematite; l o n g l i g h t dashes: i r o n c o n t e n t o f s p h a l e r i t e . A f t e r B a r t o n e t a l . (1977) .

f rom Creede, but o ther concentrations in d i f ferent samples a r e possible. However, a t this t o t a l sulfur concentra t ion (0.02 m), the pH of t h e pyr i te + hemat i t e + Fe-chlorite t r ip le point i s near 5.4, which lends credence t o t h e thermodynamic e s t ima te of pH (Fin. 7.9). "

Chemical buffering of the o re fluids--Barton et al. (1977) proposed t h a t reactions between iron-rich chlorite, hemat i te , pyrite, quar tz and water controlled t h e redox conditions for t he OH vein. A means of examining tha t suggestion arises f rom Woods et al.'s (1982) observation t h a t t he iron con ten t s of growth- banded sphalerite from the OH vein show a positive corre la t ion with t h e homogenization t empera tu res (Th) of primary fluid inclusions contained in the growth bands (Fig. 7.10). Previous investigators (Barton and Toulmin, 1964; S c o t t and Barnes, 1971) have shown t h a t t he iron content of sphaler i te in a pyrite- s a tu ra t ed system is a function solely of t empera tu re and t h e ac t iv i ty of sulfur ( the role of pressure may safe ly be neglected for these shallow deposits), based on t h e definit ive equation fo r t he iron in sphaler i te in pyrite-saturated system given below

FeS ( i n sph) + 112 S2 = FeS2 ( p y ) ( 2 )

In order t o determine if t he sys t ema t i c variation of Th with iron content is due t o a mineralogically reasonable buffer, we will examine the following reaction which could buffer t h e ac t iv i ty of sulfur in the ore-forming solution

3 d a p h n i t e + 3 K- fe ldspa r + 5 S2

= 5 h m t i t e + 5 p y r i t e + 3 h i c a

+ 9 q u a r t z + 9 H20 ( 3 )

This react ion differs f rom t h a t of t he proposed buffer of Barton et al. (1977) in t h a t t h e Fe-chlorite in this reaction (3) is daphnite (Fe5AI Si 0 (OH)g) ra ther than an aluminum-free end rnemse?. '8aphnite has a more rea l is t ic c o m ~ o s i t i o n which closelv a ~ ~ r o x i m a t e s t h a t of t h e c r e e h e chlor i te ( ~ m m o n s :Ad Larsen, 1923). However, using daphnite in the react ion requires t h a t o ther aluminum-bearing phases be considered t o balance t h e reaction. The two most logical choices for Creede a r e K-feldspar and ser ic i te , t h e two phases which Barton et al. (1977) suggest controlled the pH of t h e sys tem (discussed above).

In order to s e e if react ion (3) predicts t he observed correlation of t empera tu re with sphalerite iron content , i t is necessary t o (1) know t h e change in t h e f r e e energy of t h e react ion with t empera tu re so t h a t t he ac t iv i ty of sulfur can be es t imated, and (2) use t h a t sulfur ac t iv i ty t o ca lcula te t h e iron content of sphalerite a t t h a t temperature . Est imat ing the change in the f r e e energy of t h e react ion is l imited by the lack of f r e e energy d a t a on t h e daphnite component. However, since we a r e only trying t o predic t t h e change in the ac t iv i ty of sulfur with t empera tu re for this buffer reaction, t h e accuracy of t h e f r e e energy da ta is not a s important a s knowing how i t changes with temperature ; Hemingway etal. (1984) have recently measured the hea t capaci ty of two natura l chlorites (one Fe-rich, t h e o ther Mg-rich), and their d a t a ag ree well with Helgeson e t al.'s (1978) hea t capaci ty da ta for daphnite between 2 0 0 ~ - 3 0 0 ~ ~ . Although the heat-capacity d a t a allow us t o ca lcula te t h e change in the f r e e energy with temperature , in order t o use those d a t a i t is s t i l l necessary t o e s t ima te t h e chemical potential of daphnite a t one temperature . Therefore, we have used the e s t i m a t e of Barton e t al. (1977) of t h e chlor i te + pyr i te + hemat i t e tr iple point a t 2 5 0 ' ~ a t approximately -11.0 fo r log S2 act iv i ty and -34.2 for log O2 activity. Using these values, the chemical potential a t 2 5 0 ' ~ for t he daphnite component of t he Creede chlor i te is e s t ima ted to be -1462 kcal/mole.

Using the ac t iv i ty of sulfur e s t ima ted from t h e above thermodynamic calculation a t a given temperature , the iron con ten t of the sphalerite can be ca lcula ted using t h e following equation

l o g X(py) = 6.809 - 7340/T - 0.5 l o g a s 2

( 4

This equation which is a numerical expression of react ion (2) has t h e s a m e slope a s S c o t t and Barnes (1971) equation, but t h e in t e rcep t has been changed t o ag ree with Czarnanske's (1974) measurements on the iron content of sphaler i te in equilibrium with pyr i te +


Figure 7.9. Log aO2 - pH diagram at 2 5 0 ~ ~ and total sulfur = 0.02 molal showing the mineral stability fields pertinent to the Creede ores (modified from Barton et al., 1977). The salinity is 1 molal, with Na+/K+ = 9. Dotted lines: boundaries between aqueous sulfur species; thick solid lines: boundaries between pyrite, pyrrhotite, chlorite, and hematite; short dashes: limit of stability of chalcopyrite; long light dashes: iron content of sphalerite; medium solid lines: boundaries between kaolinite, muscovite, and potassium feldspar. Abbreviations: py = pyrite, bn = bornite, ccp = chalcopyrite, chl = chlorite.

bornite + chalcopyrite, a s discussed by Barton et al. (1977, p. 10).

In order t o directly r e l a t e t h e iron con ten t of t he sphaler i te t o the buffer, we can combine equations (2) and (3) t o ge t t he following react ion

3 d a p h n i t e + 3 K- fe ldspa r + 5 p y r i t e

+ 9 q u a r t z + 9 H20 ( 5 )

Figure 7.10 shows t h a t t h e predicted iron content between 200-280°C fairly closely ma tches t h e measured data , indicating t h a t t h e ore-forming sys tem may have been indeed buffered by react ion 13). I t is

important t o no te t h a t t h e line on Figure 7.10 was forced t o go through t h e d a t a a t 2 5 0 ' ~ by our e s t ima te of t he f r e e energy of daphnite, but t h a t t h e slope of t h e line was determined by t h e heat-capacity data. The sca t t e r in t h e d a t a in Figure 7.10, o ther than can be a t t r ibu ted t o analysis and correlation error, is presumbably due t o perturbations of t h e chemical system away f rom t h e buffered environmenf, especially a t t h e higher iron concentrations, a s discussed by Barton et al. (1977).

Sources of metals--Lead-isotope studies repor ted by Doe e t al. (1979) and unpublished d a t a of Foley (U.S.G.S., 1985) show t h a t t he lead-isotopic composition of galena f rom t h e OH, Amethyst and Bulldog Mountain vein systerns is remarkably uniform and is more radiogenic than t h a t of any volcanic rock in the San Juan Mountains, or for t h a t ma t t e r , than any Mesozoic or Cenozoic rock f rom t h e en t i r e Rocky Mountain region measured t o date. Doe e t al. (1979) conclude t h a t th is requires t h e bulk of t h e lead in t h e Creede system t o be derived by leaching of 1.4 t o 1.7 billion-year-old Precambrian rock underlying t h e district . Surprisingly, t h e lead-isotopic composition of galenas from t h e Alpha-Corsair f au l t sys tem and from the mineralized Creede Format ion a t Monon Hill, nearly on s t r ike of t h e Alpha-Corsair f au l t system, a r e less radiogenic than t h e galenas f rom t h e main par t of t h e district . These galenas appear t o have a much larger component of lead f rom t h e volcanic rocks, suggesting t h a t t he lead deposited along t h e Alpha- Corsair system was derived f rom a di f ferent , possibly more shallow, lead reservoir than t h a t of t he Amethyst-OH-Bulldog Mountain system.

Sources of sulfur--Sulfur-isotope studies a t Creede, only partially repor ted t o d a t e (Bethke e t al., 19731, indicate a complex sulfur history not ye t fully understood. Sulfur-isotopic equilibrium between sulfide minerals and reduced aqueous sulfur species was apparently closely approached. I t is clear, however, t ha t t he re was l i t t le, if any, sulfur-isotopic exchange between oxidized and reduced aqueous sulfur species in the o r e zone, and they appear t o have opera ted a s sepa ra t e isotopic (and probably chemical) sys tems in t h e upper pa r t of t h e hydrothermal system. The narrow range of sulfur-isotopic composition of t h e sulfide minerals i s i n t e rp re t ed t o r e f l ec t a deep, nearly 0 permil, sulfide reservoir, whereas t h e ext remely heavy values of 6180 and 6 3 4 ~ f rom sulfa te minerals (up t o 19 and 45 permil, respectively) require t h a t a t leas t a large par t of t h e sul fa te underwent biogenic reduction in the playa lake, t h e presumed reservoir for t he o re fluids. Some of the sul fa te was carr ied deep in to the roots of t h e Creede hydrothermal system where t empera tu res were high enough M 5 0 ° c ) fo r i t t o equilibrate with t h e wallrock s i l ica tes and t h e 0 permil sulfur reservoir. Most of t h e sul fa te did not pene t r a t e deeply enough t o a t t a i n a high enough t empera tu re t o undergo substantial isotopic exchange. This partially-to-unexchanged su l f a t e did, however, mix into t h e hydrothermal sys tem, presumably along t h e margins of t h e upwelling brine plume, t o produce the exceptionally wide range of sulfur- and oxygen- isotopic compositions measured on t h e barites.

Hydrologic Environment

The hydrologic environment is a ~ n a j o r f ac to r in t h e localization of epi thermal ore bodies. Unfortu- nately, i t is also the environment which we a r e a t present l ea s t able t o t r e a t quantitatively. However, a number of geologic, isotopic, and fluid-inclusion evidences can be used t o p lace some res t ra in ts on t h e hydrologic environment of t he Creede system, and t o develop a qual i ta t ive model which can be cornpared

Homoganization Temperature OC

Figure 7.10. Diagram showing the temperature of homogenization of fluid inclusions vs. the iron content of the host sphalerite growth zone for sample locality NJP-X on the OH vein. The line shows the predicted iron content of the sphalerite if the sulfur fugacity of the system had been buffered by the triple point - Fe-chlorite (daphnite), pyrite, hematite.

San Luis Caldera

with observed hydrologies of ac t ive geothermal systems.

Figure 7.1 1 shows a generalized hydrologic rnodel for t h e Creede mining dis t r ic t a s i t is presently understood. I t is similar in most aspects t o t h e model orginally proposed by Steven and Eaton (1975) based upon geologic grounds, and t o t h a t proposed by Barton e t al. (1977) based upon mineralogical and geochemical evidence. I t is also similar in overall aspects t o models of geothermal sys tems in the Taupo Volcanic Zone, New Zealand (cf. Henley and Ellis, 1983).

The Creede ores were deposited along the top of a saline, deeply circulating hydrothermal system, charged primarily with me teo r i c waters, t h a t displaced t h e regional ground water regime. An intrusion underlying the dis t r ic t a t a depth of 3 t o 5 km is specula ted t o have been the hea t source which provided t h e buoyancy of the brine plume. In the upper p a r t of t h e sys tem, fluid movement was fracture-controlled and nearly horizontal, from north t o south. The o re zone was overlain by a zone of fresh ground wa te r , approximately 500 me te r s thick, , which flowed southward down the regional hydraulic gradient. The base of this ground water zone was heated t o temperatures of about 160°C by hea t t ransfer f rom the underlying brine. Precipitation of t h e o res a t t h e in ter face between t h e deep brine and t h e overlying ground water appears to have resulted f rom t h e dual processes of boiling and mixing. The various evidences supporting this general hydrologic model a r e presented below.

Geologic constraints--Steven and Eaton (1975) suggested t h a t t h e circulating hydrothermal system responsible for mineralization was influenced by two major lithologic factors. The f i rs t was the location of sof t , non-welded t o poorly-welded, relatively impermeable , tuf fs along t h e top of t h e ore zone. This

NORTH SOUTH

Figure 7.11. Schematic representation of the Creede hydrothermal system. Upwelling plume (stippled pattern), outlined by the 2 0 0 ~ ~ isotherm, displaces regional ground water flow from highlands in the north to the low area of the Creede caldera moat to the south. Heat source responsible for buoyancy of plume is shown as stock beneath district (hatched pattern).


aqui tard largely blocked the upward movement of t he hydrothermal solutions and forced t h e solutions t o flow la tera l ly t o the south. The second lithologic f ac to r was t h e location of t h e permeable talus-regolith and fanglomerate deposits of the Creede Formation a t t h e southern end of t h e vein system (Figs. 7.3 and 7.11). Where c u t by t h e flow-controlling f rac tures , these coarse clastics provided an out le t t o t h e south for t h e hydrothermal solutions.

The mineralogically and geochemically based hydrologic model of Barton e t al. (1977) is nearly identical t o t h a t of Steven and Eaton (1975). Barton e t al. describe t h e system a s a f ree ly convecting hydrothermal ce l l t h a t ex t r ac t ed meta ls and sulfur f rom sources a t depth and deposited gangue and o r e minerals near t h e top of the system. They a t t r ibu te o r e deposition t o cooling and a slight pH rise due t o boiling of the hydrothermal fluid. Recondensation of t h e C 0 2 and H2S, which were strongly f rac t ionated in to t h e vapor phase during boiling, in the cooler, overlying rocks led t o t h e format ion of the intensely ser ic i t ized cap above the ore (Fig. 7.6). The development of this zone of intense a l tera t ion above t h e orebodies undoubtedly served t o increase t h e efficiency of t h e aqui tard along t h e top of the system.

Influence of topography--In addition t o t h e influence of t h e soft-tuff aquitard and ser ic i t ic a l tera t ion zone, i t is probable t h a t t h e topography a t t h e t i m e of ore-formation played an important role in maintaining a "cap" of cooler ground water above t h e deeply circulating ore-forming brine. According t o Steven and Eaton (19751, a t t he t ime of ore-formation the a r e a of t h e San Juan Mountains comprised a widespread volcanic plateau punctuated by regions of rough topography in ca ldera areas. Local relief in t h e vicinity of Creede approached 1.8 km over horizontal distances of 10 km. The low point was t h e playa lake in the moa t of t he Creede caldera, and maximum elevations were a t t a ined along a string of volcanoes located along the present Continental Divide, about 10 km nor th of t h e geographic center of t h e mining district . This high relief imposed a strong regional hydraulic gradient on t h e ground water from north t o south across t h e o r e zone. As noted by Henley (1985, this volume), such a regional hydraulic gradient will tend t o establish a la tera l ground-water flow across t h e top of a geothermal system. Such an e f f e c t has been evaluated by Hanaoka (1980) using numerical modeling techniques. Hanaokals analysis is simplified in that: ( I ) i t is calculated for pure water; (2) t h e conditions were chosen to obviate boiling; (3) t h e permeabili ty and thermal conduction-dispersion coeff ic ient were both uniform; and (4) t h e e f f e c t of pressure on viscosity was ignored. Changes in these variables will modify the quant i ta t ive aspects of Hanaoka's results, but, within reasonable bounds, a r e not likely t o a f f e c t t he general topology of his models. His calculations showed t h a t in a reas of modera t e relief, a two-tiered flow regime may be produced, consisting of a deep, hot, convecting cel l overlain by a cooler region of ground-water flow forming a "hydrologic" cap t o the hydrothermal system. Flow is parallel and l a t e ra l along t h e in ter face between t h e two regimes, and hea t is t ransferred from the deep t o t h e shallow system by conduction and dispersion.

Hanaoka's model is for moderate relief consistent with the conditions observed in a number of geothermal systems (Healy and Hochstein, 1973; Healy, 1975; Ellis and Mahon, 1977; Hedenquist , 1983). I t is also tantalizingly consistent with the evidence for t he in teract ion between a hot, deep, saline fluid, and a cooler, overlying fresh ground water a t Creede.

Isotopic constraints--Light-stable-isotope studies have shown t h a t in addition t o the deep, saline fluids and t h e shallow groundwaters discussed above, fluids from a third, isotopically distinct, probably magmatic, source were involved in t h e Creede mineralization. The evidence for t h e episodic introduction of magmat i c waters in to t h e Creede hydrothermal system comes f rom carbon- and oxygen-isotopic studies of t he vein carbonates (Bethke and Rye, 1979). These studies indicate t h a t both t h e ear ly (A-stage) rhodochrosite and t h e younger (C-stage) siderite-manganosiderite were deposited from fluids t h a t equilibrated with sil icates a t very high t empera tu res (presumably magmat i c fluids). This supposition was based primarily on t h e large 6180 values obtained, but i t is also consistent with the carbon-isotopic composition and with the hydrogen-isotopic composition of t h e inclusion fluids in a few rhodochrosite samples. The in terpre ta t ion remains open t o question, however, because i t requires e i ther t h a t t he vein system was occupied by magmat i c fluids a t t w o di f ferent t imes , separa ted by a period when i t was occupied by t h e me teo r i c wa te r s f rom t h e lake sediment reservoir, or t ha t C02 .g iven off from t h e magma did not exchange oxygen w ~ t h t h e me teo r i c waters t h a t filled the vein system. Neither seems a reasonable proposition. Carbonate minerals in many hydrothermal o r e deposits appear t o have been deposited from fluids substantially heavier o r l ighter in 6180 than those from which other apparently coeval minerals were deposited (R.O. Rye, U.S.G.S., personal communication, 1985). Perhaps we do not fully understand t h e isotope sys temat ics of C 0 2 in hydrothermal systems.

Based on t h e hydrogen- and/or oxygen-isotopic composition of the a l tera t ion minerals, and chlor i te and t h e inclusion fluids in sphalerite, Bethke and Rye (1979) postulated t h a t t he deep, saline fluids responsible for ore deposition originated a s pore- wa te r s in t h e Creede Format ion which accumulated in the evaporative, closed-basin playa lake in t h e moa t of t he Creede caldera. Such an in terpre ta t ion is consistent with the unusually high salinit ies of t he Creede o re fluids, and with sulfur- and oxygen-isotopic s tudies on bar i te (R. Rye, P. Barton and P. Bethke, U.S.G.S., unpublished data , 1985) which suggest t h a t most of t h e sul fa te in t h e Creede system must have undergone considerable isotopic evolution in the playa lake sediments.

The recent work of Foley e t al. (1982) demonstra ted t h e exis tence of a zone of heated, f resh me teo r i c water overlying the o re zone. They showed t h a t fluids in primary inclusions in qua r t z were similar in salinity, homogenization temperature , and hydrogen-isotopic composition (and therefore , presumably, source) t o primary inclusions in sphalerite. This in terpre ta t ion was contrary t o t h a t originally proposed by Bethke and Rye (1979), which suggested t h a t t he quar tz was deposited f rom fluids

t h a t originated a s me teo r i c waters in the high country nor th of t h e district . Very painstaking sampling and analysis showed t h a t the hydrogen-isotopic analyses of fluid inclusions in quar tz repor ted by Bethke and Rye (1979) were biased by fluids released f rom pseudosecondary inclusions during analysis. These pseudosecondary inclusions homogenized over a similar but slightly lower t empera tu re range a s did t h e primaries, but contained fresh waters whose hydrogen isotopic composition was a t l ea s t 30 permil l ighter than t h a t of t h e brine in t h e primary inclusions. Foley et al. (1982) proposed t h a t these light, f resh waters were unevolved me teo r i c waters t h a t const i tu ted the regional ground waters. These fluids overlay t h e o re zone, but episodically en te red i t due t o hydrologic fluctuations during vein filling. Presumably, t h e thermal shock of these heated, but somewhat cooler ground waters , caused f r ac tu res in growth-strained qua r t z crys ta ls which, on rehealing, t rapped some of the fluids.

The lead-isotope s tudies by Doe e t al. (1979) and unpublished d a t a of Foley (U.S.G.S., 1985), discussed in t h e section on sources of metals, also provide evidence t h a t t h e Creede hydrothermal system ci rcula ted t o depths of several kilometers. Their d a t a suggest t h a t t h e bulk of t he lead in t h e Creede system was derived by se lect ive leaching of radiogenic lead f rom t h e 1.4 t o 1.7 billiowyear-old Precambrian rock underlying the district . The depth t o t h e Precambrian a t Creede (and hence t h e minimum depth of hydrothermal circulation) can only be es t imated, but must be a t l ea s t 2.5-3.0 km below the ore zone (T. Steven, U.S.G.S., wr i t ten communication, 1982).

Boiling and Mixing in the Ore Zone

The most recent modification t o the in terpre ta t ion of the hydrology is t h e evidence for mixing of the hydrothermal fluid with overlying ground water developed by Hayba (1984) for t h e OH vein and by Robinson and Norman (1984) fo r t h e southern Amethyst vein. Although boiling has been documented in t h e upper levels of t he OH vein (Roedder, 1970; Woods et al., 19821, i t appears t h a t mixing was responsible for a t l ea s t some of the vein mineral deposition. These two processes, boiling and mixing, a r e each important mechanisms of mineral deposition; evidences for their roles in t h e Creede hydrothermal system will be discussed below.

Boiling--One of the major f ac to r s influencing the hydrologic in terpre ta t ion of t h e d is t r ic t is t he f a c t t h a t boiling occurred during a t l ea s t some of the depositional history. As noted ear l ier , t h e r e a r e very few deposits exhibiting unequivocal evidence t h a t boiling occurred during precious-metal deposition. A t Creede, t he re is good evidence for boiling, but i t appears t o have taken place during s t age D, a base- meta l s tage , ra ther than during the precious-metal s t age B (Table 7.4). The case for boiling a t Creede is based primarily on two lines of evidence: (1) fluid inclusions, and (2) t h e distributions of vein adularia and ser ic i t ic alteration.

Over 25 years ago, Roedder (1960) discovered some coeval fluid inclusions in D-stage qua r t z t h a t had widely varying IiquidJvapor ratios, and thus a wide range of homogenization temperatures , many over

280°C (about t h e maximum tempera tu re for Creede mineral deposition based on d a t a f rom over 2,500 fluid inclusion measurements). H e in terpre ted these inclusions a s having resulted f rom t h e trapping of varying proportions of liquid and vapor from a heterogeneous, two-phase system. This was t h e f i rs t such demonstration of boiling in any ore deposit. His in terpre ta t ion is fu r the r bols tered by the presence of "empty" or "steam" inclusions in some of these same qua r t z samples (e.g., Roedder 1970, Fig. 7.71, and by the occasional occurrence of o re s t a l ac t i t e s (indicative of growth in a two-phase regime) protruding in to open vugs along the OH and Bulldog Mountain vein systems.

T o date , evidence for boiling has been found a t six locali t ies along t h e top of t h e OH vein (Fig. 7.6) in D-stage quartz. I t is notable t h a t Robinson and Norman (1984) found no d i r ec t evidence of boiling in qua r t z samples f rom t h e southern Amethyst vein although they carefully looked for indications of boiling. No such evidence has ye t been found in t h e l imi ted number of fluid-inclusion measurements on mater ia l from t h e Bulldog Mountain vein system, but t he mater ia l so f a r examined has come from a reas on t h a t vein which l ie t o t h e south of t h e a rea on t h e OH from which boiling has been documented. The implication is t h a t boiling was primarily confined t o the northern par t of t h e d is t r ic t , a t leas t in t h e la ter s tages of o r e deposition during which most of t h e mater ia l studied t o d a t e was deposited. This surmise is consistent with t h e extensive evidence for mixing (discussed below) of t h e o re fluids with cooler overlying fresh wa te r s along t h e top of t h e system. To date , fluid-inclusion evidence fo r boiling a t Creede has been found only in qua r t z crystals, even though i t has been sought carefully in sphaler i te and fluorite. I t i s possible t h a t boiling a lso occurred during t h e "B-stage" deposition of the precious-metal ores, but evidence for boiling during this s taged has been carefully looked for and has not been found. The fine-grained nature of B-Stage mater ia l s t age makes recognition of fluid- inclusion evidence for boiling difficult , but t h e lower homogenization t empera tu res for B-Stage inclusions than for D-Stage inclusions is consistent with a lack of boiling.

Also shown on Figure 7.6 is t he distribution of vein adularia and in tense ser ic i t ic alteration. Adularia occurs mainly below t h e zone of boiling documented by fluid inclusions, while the zone of intense ser ic i t ic a l tera t ion, which caps the OH vein, is above t h e zone of documented boiling. The ser ic i t ic cap, which is spectacular in i t s contras t and sharpness of con tac t with t h e unaltered wallrock below ( the productive pa r t of t he vein), may have resulted f rom the recondensa- t ion of volati les in t h e overlying, probably fresh waters and from the subsequent hydrolysis reactions. A t Broadlands, adular ia and ca l c i t e deposition has been re la ted t o the r i se in pH of t h e fluid on boiling, due t o t h e s t rong f rac t ionat ion of acid-forming volati les such a s C 0 2 and H2S in to the vapor phase (Browne and Ellis, 1970). This pa t t e rn of in tense ser ic i t ic a l tera t ion overlying a zone of vein adularia (+ carbonate in some mining districts) provides indirect, but potentially useful evidence for boiling in epi thermal systems. I t may be a particularly useful indicator of boiling in deposits where no mater ia l

200 220 240 260 280

HOMOGENIZATION TEMPERATURE PC)

W

0 K g 20,000 I U

water

HEAT CONTENT Jlg

Figure 7.12. a). Plot of homogenization temperature (Th) vs. freezing temperature (Tf) for 221 primary inclusions in a 5-cm band of zoned sphalerite from Creede, Colorado, after Reodder (1977). The numbered areas include all data points from each of the 20 zones sampled, numbered in sequence from w n e 1 (earliest) to zone 20 (latest). The number of inclusions in each of the areas outlined are as follows: 1(2), 2(18), 3(1), 4(21), 5(27), 6(9), 7(4), 8(9), 9(8), 10(4), 11(15), 12(2), 13(7), 14(12), 15(14), 16(11), 17(4), 18(32), 19(13), 20(8). b). Plot of heat content vs. chloride content for primary inclusions (solid triangles) replotted from Figure 7.12a. Data points for surface water and steam are also shown. Tick marks for upper and lower temperature scale are offset because of the effect of chloride content.

a d e q u a t e f o r fluid-inclusion s t u d i e s i s ava i lab le , o r w h e r e t h e fluid-inclusion e v i d e n c e f o r boiling i s ambiguous.

T h e v i r t u a l l a c k of c a l c i t e in t h e C r e e d e d i s t r i c t c a n b e expla ined by t h e low-calcium c o n t e n t of hos t r o c k s resu l t ing f r o m t h e e a r l i e r potassiuln m e t a s o m a t i s m e v e n t , a n d t o t h e l o w - C 0 2 c o n t e n t of t h e f luids during m o s t of t h e depos i t iona l history. W e t l a u f e r (1977) h a s shown t h a t t h e f lu ids responsible

f o r deposi t ion of t h e e a r l y (A-Stage = S t a g e I) Mn-Fe c a r b o n a t e s i n t h e l o w e r p a r t s of t h e s o u t h e r n por t ions of t h e vein s y s t e m s w e r e s imi la r in t h e i r t h e r m a l h i s tory a n d c h e m i s t r y t o t h e l a t e r f luids which depos i ted t h e bulk of t h e base- a n d prec ious-meta l minera l iza t ion e x c e p t t h a t t h e carbonate -depos i t ing f luids had a h igher CO-, c o n t e n t . I t i s n o t unreasonable t o suppose t h a t t h e e a r l y Mn-Fe c a r b o n a t e s in t h e C r e e d e vein sys te rns p lay t h e r o l e of c a l c i t e in t h e Broadlands g e o t h e r m a l s y s t e m a n d r e c o r d a period of boiling of re la t ive ly gas-r ich f lu ids e a r l y in t h e h i s tory of vein filling.

Mixing--Roedder (1977) showed t h a t t h e r e was a s y s t e m a t i c r e l a t i o n b e t w e e n t e m p e r a t u r e a n d sa l in i ty f o r f luid inclusions f r o m t h e O H vein. H e d o c u m e n t e d t h i s re la t ionsh ip b y a d e t a i l e d growth-zone by growth- z o n e s tudy of a s ingle, l a r g e s p h a l e r i t e c r y s t a l f r o m a single loca l i ty (NJP-X) o n t h e OH vein. T w e n t y d i s t i n c t g r o w t h z o n e s w e r e d e f i n e d in t h i s c r y s t a l , and homogeniza t ion a n d f r e e z i n g t e m p e r a t u r e s w e r e m e a s u r e d o n s e t s of inc lus ions wi th in e a c h zone. T h e r e s u l t s of h i s pa ins tak ing s t u d y a r e shown in F i g u r e 7.12a. S imi la r r e s u l t s w e r e o b t a i n e d f o r s e v e r a l d i f f e r e n t l o c a l i t i e s o n t h e O H vein (Woods e t al., 1982). T h e r e a r e m a n y impl ica t ions of t h e s e s y s t e m a t i c s , p e r h a p s t h e rnost i m p o r t a n t f o r our p r e s e n t purposes be ing t h e mix ing of f lu ids of d i f f e r e n t t e m p e r a t u r e s a n d sal ini t ies . Truesde l l a n d Fourn ie r (1975) h a v e shown t h a t p lo t s of h e a t c o n t e n t (en tha lpy) a g a i n s t ch lor ide c o n t e n t a r e v e r y usefu l in e v a l u a t i n g t h e re la t ionsh ip b e t w e e n f luids of d i f f e r e n t t e m p e r a t u r e s a n d compos i t ion in g e o t h e r m a l a reas . Both e n t h a l p y a n d c h l o r i d e c o n t e n t a r e addi t ive q u a n t i t i e s s o t h a t t r a j e c t o r i e s f o r boiling a n d mix ing a r e l inear o n s u c h plots. Roedder ' s d a t a a r e r e p l o t t e d a s en tha lpy-ch lor ide d iagrarns in F i g u r e 7.12b. I t c a n b e s e e n f r o m t h i s f i g u r e t h a t t h e s y s t e r n a t i c re la t ionsh ip b e t w e e n t e m p e r a t u r e a n d sa l in i ty c a n b e expla ined by mix ing of sa l ine , h i g h - t e m p e r a t u r e w a t e r s (from z o n e s 8 a n d 9 o n F i g u r e 7.12a) wi th f r e s h w a t e r s h e a t e d t o a t e m p e r a t u r e of a b o u t 1 6 0 ' ~ . I t i s possible t h a t w a t e r s f r o m z o n e s 8 a n d 9 a r e r e l a t e d t o e a c h o t h e r in t h a t a s m a l l a m o u n t of boiling of z o n e 8 w a t e r s would yield t h e s l igh t ly m o r e sa l ine , lower t e m p e r a t u r e w a t e r of z o n e 9. Enthalpy- o r chloride- conserva t ion c a l c u l a t i o n s i n d i c a t e t h a t if t h e s t e a m i s s e p a r a t e d f r o m t h e fluid, on ly 6 wt.-% of t h e f luid needs t o b e c o n v e r t e d t o s t e a m t o produce w a t e r 9 f r o m w a t e r 8. O n t h e o t h e r hand, in o r d e r t o produce t h e l o w e s t - t e m p e r a t u r e , l eas t - sa l ine f luid ( w a t e r 15), w a t e r 8 would h a v e t o b e m i x e d wi th m o r e t h a t i t s equiva len t w e i g h t of f r e s h w a t e r a t 1 6 0 ' ~ .

Roedder 's d o c u m e n t a t i o n of t h e a b r u p t t e m p o r a l var ia t ions i n t h e o r e f luid a l o n g t h e OH s t r u c t u r e w a s a n i m p o r t a n t t i m e c o n s t r a i n t f o r a l a t e r s t u d y on t h e s p a t i a l v a r i a t i o n s i n t h e o r e fluid. Using a d i s t i n c t i v e g r o w t h z o n e i n s p h a l e r i t e a s a t ime- l ine t h r o u g h o u t t h e O H vein, H a y b a (1984) d o c u m e n t e d a progress ive d e c r e a s e in b o t h t e m p e r a t u r e and sa l in i ty f r o m t h e nor thern , basa l e n d of t h e ve in t o loca l i t i es 200 m e t e r s higher and 1000 m e t e r s f u r t h e r south. T h e s e t e m p e r a t u r e a n d s a l i n i t y g r a d i e n t s a r e i n t e r p r e t e d as t h e progress ive mix ing of d e e p e r , sa l ine h y d r o t h e r m a l f luids w i t h over ly ing , d i l u t e ground w a t e r s t h a t h a v e b e e n p r e h e a t e d t o a p p r o x i m a t e l y 1 6 0 ' ~ . Independent ,

isotopic evidence for t h e presence of a dilute ground water in t h e OH vein has been documented by Foley e t al. (1982) (discussed earlier). I t is in teres t ing t o no te t h a t t h e e s t ima ted t empera tu re of 1 6 0 ' ~ for t h e dilute ground wa te r s i s within the 100' t o 1 8 0 ~ ~ t empera tu re range e s t ima ted for t he diluent in most New Zealand geothermal sys tems (Hedenquist and Reid, 1984).

The fluid-inclusion s tudies of Robinson (1981; Robinson and Norman 1984) also indicate t h a t t h e deep hydrothermal solutions mixed with shallow ground water in t h e southern Amethyst vein. Due t o t h e f a c t t h a t t he i r fluid-inclusion s tudy was done on quar tz , i t was impossible for them t o distinguish growth zones or make deta i led t i m e correlations. Instead, a more general approach was t aken and the inclusions f rom one s t age of qua r t z deposition were measured on samples covering a ver t ica l range of 336 meters. Although a t e a c h sampled elevation the re is a large range in both t h e t empera tu re and salinity in t h a t s tage , t he re i s a general decrease in both t empera tu re and salinity with increasing elevation, which they a t t r ibu te t o mixing.

Additional evidence fo r mixing comes f rom the district-wide mineral zonation of t h e sulfide-rich OH vein in t h e nor th t o t h e barite-rich Bulldog vein in t h e south. In Figure 7.13 t h e solubility of bar i te is contoured on a temperature-salinity diagram and mixing and boiling t ra jec tor ies relevant t o Creede a r e superimposed. I t can be seen t h a t bar i te solubility changes relatively l i t t l e a t t h e high temperatures and salinit ies appropr ia te for input fluids in the northern OH vein, and drops significantly only a t salinit ies below about 6 wt.-% NaCl. equivalent (Plumlee and Hayba, 1985). (Note: in sulfate-rich solutions above pH 5, changes in pH have no e f f e c t on bar i te solubility). The topology of Figure 7.13 suggests t h a t fluid mixing was t h e depositional mechanism for bar i te a t Creede. Most mixing paths (decreasing t empera tu re and salinity) cross solubility contours while boiling paths a r e parallel t o them. Thus, only a f t e r t h e hot, saline fluids, which rose in t h e northern par ts of t h e d is t r ic t , were diluted significantly did they deposit t h e large quant i t ies of bar i te seen in t h e southern and upper par ts of t h e district . The silver content of t h e Creede o res i s also higher in t h e southern and upper par ts of t h e d is t r ic t (Barton et al., 1977), suggesting t h a t mixing was an important mechanism of silver deposition a t Creede.

Rela t ive importance of boiling and mixing--It was pointed ou t a t t h e beginning of this sect ion t h a t both boiling and mixing a r e important processes leading t o t h e deposition of base- and precious-metal ores in epi thermal systems. Drummond and Ohmoto (19851, Henley (1985, th is volume) and Reed and Spycher (1985, this volume) have a l l emphasized the ef f ic iency of boiling in t h e precipitation of both ore and gangue minerals in epi thermal environments. Barton e t al. (1977) specifically re la ted t h e precipitation of t h e Creede o res t o cooling and pH changes occasioned by boiling. The evidence from t h e recent s tudies of Creede c i t ed above, however, would imply t h a t mixing, not boiling, was t h e immediate cause of most, if not

100 140 180 220 260 300

TEMPERATURE ("C)

Figure 7.13. Diagram showing the solubility of barite contoured on a temperature vs. salinity plot. Barite solubility was calculated at pH = 5.5, mS = 0.02, sQ,/Sred = 10, .mNa/my = 9. Arrows mark boiling and mlxing trajectories.

a l l o r e depostion a t Creede. This evidence is summarized below: I. Fluid-inclusion evidence for boiling has been found

only on t h e northern half of t h e OH vein, and only in qua r t z deposited in t h e la tes t , silver-poor, s t age of mineral deposition on t h e OH.

2. Fluid-inclusion evidence for substantial amounts of mixing, increasing upward and t o t h e south, have been found along the OH and southern Amethyst vein systems.

3. The deposition of large amounts of bar i te in t h e southern par ts of t h e d is t r ic t and i t s absence in t h e northern par ts i s consistent with a mixing model, no t with a boiling model of bar i te deposition.

4. On t h e OH vein, t he maximum tempera tu re measured on B-Stage mater ia l ( the principal silver- bearing s t age a t Creede) is 241°C, at l eas t 30' less than t h e maximum tempera tu re measured f rom D- S tage material . To produce boiling a t t h a t t empera tu re would require t h a t t h e water t ab le be lower by about 200 me te r s during B-Stage than i t was during D-Stage.

5. In a system dominated by la tera l flow (i.e., one t h a t allows for only minor changes in pressure), i t is virtually impossible for a fluid t o boil a f t e r i t s t empera tu re has been lowered by mixing.

The widespread and intense ser ic i t ic a l t e ra t ion t h a t caps the Creede orebodies, t he widespread occurrence of vein adularia in t h e sys tem, and t h e deposition of massive amounts of Mn-Fe carbonates in t h e deep par ts of t he southern portions of t h e vein sys tems all a rgue t h a t very substantial amounts of boiling took place during vein filling. None of these f ea tu res can, however, be corre la ted with t h e major periods of base- and precious-metal deposition. We conclude, therefore , t h a t although boiling can be well documented a t Creede by several cr i ter ia , t he re is no

evidence t h a t boiling ( a t leas t in t h e ore zone) was re la ted to base- o r precious-metal deposition. The evidence fo r mixing of t h e deep, saline, o r e fluid with overlying f resh ground water during sulfide deposition, on t h e o the r hand, is overwhelming. As pointed ou t in the f i rs t sec t ion of this chapter , Creede is not unique among t h e Adularia-Sericite-type deposits in exhibiting a lack of evidence for boiling t ied t o me ta l deposition. Only a t Colqui is t h e evidence compelling t h a t precious-metal deposition resulted from boiling of the ore fluid (Kamilli and Ohmoto, 1977).

Summary of Creede mineralization

On both geologic and geochemical grounds, i t has been proposed t h a t t h e Creede o res were deposited along t h e top of a deeply circulating hydrothermal system a t t h e in t e r f ace of t h a t system with t h e overlying ground water. The concept is i l lustrated in Figure 7.11. The hea t source driving t h e hydrothermal circulation is speculated t o have been an intrusion underlying t h e dis t r ic t a t a depth of 3 t o 5 kilometers. This intrusion may have been re la ted t o t h e quartz-lati te, r ing-fracture volcanism of t h e Creede caldera cycle, or may have been a qua r t z porphyry re la ted t o t h e la ter bimodal basalt-rhyolite volcanism. The o re fluids were dominated by me teo r i c waters, whose isotopic composition and salinity evolved by processes of evaporation and diagenesis in the playa lake in the moa t of t he Creede caldera. The episodic introduction of magmat ic fluids in to t h e circulating system is suggested by the isotopic composition of rhodochrosite and s ider i te in the veins. The bulk of t he lead in the Creede ores, and therefore , presumably, most of t he o ther metals, appears t o have been leached from Precambrian basement rocks a t depth. The sulfur in t h e Creede o res may have come f rom several sources and i t s isotopic composition documents a complex mixing and exchange history no t now satisfactorily understood. Most of t h e sul fa te in t h e Creede system appears to have undergone considerable isotopic evolution in t h e playa lake sediments, but t h e sulfide sulfur appears t o have been buffered by, and perhaps derived f rom, a large reservoir of magmat ic sulfur a t depth. Precipitation of t h e ores along t h e top of t h e system appears t o have resulted from the dual processes of boiling and mixing, but mixing appears t o have been quantitatively the more important mechanism of sulfide deposition. The intense, mixed-layer i l l i te /smect i te a l tera t ion cap is in terpre ted t o have been generated by condensation along t h e top of the system of acid volati les distilled off t he deeply circulating o re fluids.

The model includes substantial amounts of fluid- rock in teract ion and chemical and isotopic exchange in the deeper, higher t empera tu re pa r t s of t he system. In t h e ore zone, however, isotopic exchange between oxidized and reduced aqueous sulfur species was minimal and t h e aqueous sul fa te apparently did not exchange oxygen with the ore fluid nor was the re significant oxygen isotopic exchange between t h e o r e fluid and t h e unaltered wallrocks. The pH, the ac t iv i ty of S2 gas and redox s t a t e of t he o re fluids were buffered in the o re zone by reaction with a vein-filling

assemblage comprising: Fe-chlorite + hemat i t e + pyr i te + K-feldspar + ser ic i te (or mixed-layer illite- s m e c t i t e clays). Throughout most of t h e history of vein filling, t he redox s t a t e of the o r e fluid was t h a t of t h e t r ip le point: Fe-chlorite + pyr i te + hematite. During the early par t of sulfide deposition, however, numerous excursions t o substantially lower oxidation (and sulfidation) s t a t e s occurred. The excursions have been in terpre ted t o have resulted from episodic introduction of magmat ic emanat ions in to t h e circulating sys tem or t o reaction with fer rous s i l ica tes episodically exposed t o t h e circulating fluids deep in the sys tem through t ec ton ic adjustments. Alternatively, t h e relatively oxidized s t a t e of t he o r e fluid during most of t h e ore deposition cycle could be due t o mixing of a deep, reduced fluid with surrounding and overlying oxidized groundwater prior t o enter ing t h e o re zone, t he excursions resulting f rom lesser amounts of mixing. Present evidence does no t allow us t o choose between these two alternatives.

THE ACID-SULFATE ENVIRONMENT: SUMMITVILLE AS AN EXAMPLE

The recent thesis by Stoffregen (1985) combined with t h e ear l ier work by Steven and R a t t 6 (1960a) and several o ther studies (Patton, 1917; Mehnert et al., 1973; Lipman, 1975; Perkins and Nieman, 1983) make t h e Summitvil le mining district , Colorado, t h e best documented Acid-Sulfate-type deposit and t h e logical choice for i l lustrating the character is t ics of Acid- Sulfate-type epi thermal systems. Many of the in terpre ta t ions on Acid-Sulfate-type deposits have been made in l ight of t he important insights gained f rom t h e studies done a t Goldfield, Nevada (Ashley, 1974; Ransome, 1909). Even so, i t should be noted t h a t t h e observational base for this t ype of deposit is st i l l much smaller t han t h a t for t h e Adularia-Sericite deposits. Since our experience in t h e Summitvil le d is t r ic t is l imited, most of t h e geologic and mineralogic character is t ics discussed below, excep t where noted, a r e based on t h e work done by Steven and R a t t 6 (1960a) and by Stoffregen (1985).

As was t h e case in using Creede a s t h e exemplar fo r Adularia-Sericite-type deposits, Summitvil le has some character is t ics which a r e not representa t ive of all of t he Acid-Sulfate-type deposits. In particular, Julcani, Lake C i ty 11, and Red Mountain a r e silver- rich, contain relatively more base me ta l s (particularly lead and zinc) and appear t o have formed a t greater paleodepths. Although these and o the r d i f ferences exist , the similarit ies in mineralogy, a l tera t ion, t ec ton ic sett ing, and t iming of o r e deposition re la t ive t o host emplacement among Acid-Sulfate-type deposits a r e striking, and indicate mineralization in a distinct geothermal environment. Ashley (1982) has summarized t h e observations on a number of deposits of this type into a particularly useful occurrence model which provides fur ther docurnentation on t h e character is i t ics of Acid-Sulfate-type deposits.

Geologic and Mineralogic Character is t ics

Volcanic history--The Summitvil le mining

CHAPTER 7

Figure 7.14. Generalized geology of the Platoro and Summitville calderas, modified from Steven and Lipman (1976). Location of Summitville mining district is shown by pick-and-hammer. Control is moderate to good where boundaries are shown by solid symbols. A-A' marks location of cross- section shown in Figure 7.15.

dis t r ic t is located a t a n elevation of about 12,000 f e e t on t h e nor thwest edge of t h e Pla toro caldera and t h e younger nes ted Summitvil le ca ldera in the eas tern San Juan Caldera complex, Colorado (Liprnan, 19751 (Fig. 7.1). The quar tz l a t i t e porphyry a t South Mountain, which hosts t h e deposit , is a lava dome emplaced 22.8 rn.y ago (Mehnert e t al., 1973) along the western margin of t he Summitvil le ca ldera a t i t s intersection with the Pass Creek-Elwood Creek-Platoro faul t zone, a major s t ruc tu re cu t t ing across the cen te r of t he Platoro-Summitville ca lderas (Fig. 7.14). Drilling has confirmed Steven and Rat tg ' s (1960a) suggestion tha t t he dome is funnel-shaped in cross sect ion with a narrow intrusive pipe a t depth which f lares out near the surface (Fig. 7.15). The quar tz l a t i t e is character ized by 2 t o 8 c m K-feldspar phenocrysts in an aphanitic, greenish groundmass; plagioclase phenocrysts a r e common but a r e finer grained than the K-feldspar. Also typical a r e locally resorbed quar tz eyes (0.2 t o 2 cm), euhedral bioti te books (1 t o 2 cm), and a p a t i t e up t o 0.5 crn long (Stoffregen, 1985). The qua r t z l a t i t e a t South Mountain is bordered on the north, ea s t , and south by t h e approximately 29 m.y. old Summitvil le Andesite which filled t h e Pla toro caldera a f t e r collapse and on t h e west by the rhyodacite of Park Creek (about 28 t o 26 m.y.), which comprises lava-dome complexes erupted around the northwest margin of t h e Surnmitville caldera a f t e r i t s final collapse (Liprnan, 1975). The rhyolite of Cropsy Mountain, a coarsely porphyrit ic lava flow, locally overlies the qua r t z l a t i t e south of t he district . In

2 3 z = C .- South Mountain volcanic dome 13 z>/------------ --. I) ,----. ?: '.-

7

ml 1 \ Talus-breccia I.---.

\ A4

Cropsy Mountain Rhyolite 20.2 m.y

0 Coarse porphyry

E X P L A N A T I O N

j,",r:l South Mountain Summitville Andesite Quartz Latite 22.9 m.y. -29 m.y.

Quartz-alunite 22.3 m.y. 0 Altered area

Figure 7.15. Geologic cross-section of the restored South Mountain volcanic dome, modified from Steven and ~attb (1960a). Fault is shown with heavy line; contacts shown with thin line; both are dashed where approximate.


con t r a s t t o t h e older units, t h e rhyolite of Cropsy Mountain is everywhere unaltered. I t has been da t ed a t 20.2 i0.8 m.y. (Mehner t e t al., 1973) and a t 18.5 i1.2 m.y. (Perkins and Nieman, 1983). The a g e of mineraliza t ion is t he re fo re b racke ted between 22.8 and 20.2 m.y. in ag reemen t wi th Mehnert e t al.'s (1973) d a t e of 22.3 m.y. on hydrothermal a luni te from Summitville.

Ore deposits--The o r e bodies a r e localized along t h e southwestern margin of t h e coarsely porphyrit ic co re of t h e dome, just nor th of t h e northwest-trending faul t zone a t South Mountain (Steven and Ra t t6 , 1960a). The o re occurs in a ser ies of irregular pipes and vein-like masses of qua r t z and a luni te t h a t formed largely by r ep l acemen t of t h e original qua r t z la t i te . Significant mineralization is confined t o a ver t ica l in terval of approximate ly 400 meters , and t h e surface outcrop of t h e minera l ized zone can be circumscribed by an elipse with axes of 1.5 and 1.0 kilometers ( R a t t b and Steven, 1960a). Production through 1947 was 113,000 o z of Au, 240,OO o z of Ag, and 433,000 lbs of C u (Vanderwilt, 1947). Productions since 1947 has been insignificant.

Wallrock alteration--The a l tera t ion in t he upper pa r t of t h e deposit shows a well-defined pa t t e rn (Fig. 7.16) from a n in tense zone of ac id leaching, r e f e r r ed t o a s "vuggy sil ica alteration", surrounded by quar tz- alunite a l te ra t ion grading outward in to quar tz- kaolinite. Fu r the r out , t h e r e is generally a n abrupt change in to a n illitic* zone which grades ou t t o a montmoril lonite zone. The width of each of t h e zones is highly variable.

*Stoffregen (1985) has used t h e t e r m il l i te t o r e f e r t o a fine-grained ( < 2 micron) phyllosilicate with a 10- angstrom basal spacing, which does not expand on glycolation, and conta ins less t han 5% s m e c t i t e layers.

The vuggy sil ica a l t e r a t ion is in terpre ted t o be t h e result of t h e ac id dissolution of all t h e primary rock-forming minera ls excep t quartz. Most of t he o r e occurs in this very permeable zone, which is character ized by l a rge voids due t o t h e removal of t h e K-feldspar phenocrysts. A t t h e surf a c e t h e vuggy silica is qui te extensive, but below a depth of about 1000 f ee t , i t becomes much more res t r ic ted (Fig. 7.17). In t h e upper pa r t of t h e deposit , t he surrounding

.quartz-alunite zone is up t o 5 0 f e e t wide. Alunite occurs a s pseudomorphous r ep l acemen t s of K-feldspar phenocrysts and a s m a t t e d agg rega te s replacing t h e groundmass. With decreas ing elevation (between 11,500 and 11,000 fee t ) , a luni te becomes insignificant and quartz-kaolinite r a the r than quartz-alunite surrounds t h e vuggy sil ica "vein". The t ex tu re of th is deeper kaolinite indica tes t h a t i t has complete ly replaced pre-existing alunite. Except fo r th is deep zone, t he quartz-kaolinite zone is generally thinner (and is locally absent ) t han t h e quartz-alunite zone which i t surrounds (Stoffregen, 1985).

Outside of t h e quartz-kaolinite zone, t h e t e x t u r e of t h e a l tera t ion shows a marked change f rom ha rd and br i t t le rock, due t o a si l icif ied matr ix , t o a sof t , incompetent rock, due t o a rnatrix of clay minerals.

Unaltered Quartz- quartz Montrnorillonite-chlorite Illite-kaolinite alunite latite zone zone zone

A + v A

Y h

chloride-rich ~ontmdrillonite Illitid rock rock rich rock / Mineralized vuggy

quartz rock

Figure 7.16. Diagram showing hydrothermal alteration pattern in the Summitville district adapted from Steven and ~ a t t s (1960a).

Figure 7.17. Schematic cross section of the alteration patterns and mineral zonation of the Summitville deposit, modified from Stoffregen (1985). The clay alteration zones refer to zones 3-6 in Figure 7.16. CV - covellite, luzonite, enargite, pyrite, marcasite, chalcopyrite, trace sphalerite, sulfur, and gold assemblage; TN - chalcopyrite, tennantite, pyrite, plus minor sphalerite and trace galena assemblage.

The c lay envelope around t h e vuggy silica and quartz- alunite-kaolinite a l t e r a t ions is generally a t leas t 100 f e e t wide. According t o Stoffregen (1985), i l l i te predominates n e a r t h e quartz-kaolinite zone, but montmoril lonite becomes predominant fur ther away f rom t h e vein. Kaolinite is present throughout t h e c lay zone, but i t dec reases in both abundance and

crystall inity away f rom t h e quartz-kaolinite zone. The i l l i te becomes coarser grained with depth. [Mixed- layered i l l i te /smect i te clays (having g rea te r than 5% smec t i t e ) a r e generally no t present, occurring only locally in t h e centra l portion of the deposit where they appear t o b e related t o a post-mineral intrusion.

Mineralogy--Stoffregen (1985) describes the mineralogy of Summitvil le in t e rms of th ree main stages: main, la te , and supergene. Mineralization from t h e s e th ree s tages was la ter than much of t h e acid-sulfate alteration, a s evidenced by the presence of o r e minerals in cross-cutting f r ac tu res and voids in vuggy silica. In addition, a very minor, ear ly s t age of mineralization is also present a s 10 t o 40 micron-sized inclusions in pyrite grains of e i ther pyrrhotite or chalcopyr i te + bornite + digenite. The significance of this minor, early s t age is not clear.

The cha rac te r of t h e main-stage mineralization changes with depth in t h e deposit (Fig. 7.17). In t h e deeper pa r t s of t h e deposit a chalcopyrite-tennantite assemblage predominates. Tennantite usually contains only minor antimony and t r ace amounts of silver, although rare ly i t may contain up t o 2.0 wt.-% silver. Pyr i te plus minor, low-iron (usually 5 1.0 wt.-%) sphaler i te and a t r a c e of galena a r e par t of this deep assemblage. In t h e upper pa r t of t h e deposit, which contains t h e economically significant precious-metal mineralization, t he main-stage mineral assemblage consists of covellite, luzonite, and enargi te with pyrite, marcas i te , locally sulfur, t r a c e amounts of sphalcrite, and gold. Chalcopyrite is also present, but i t decreases in abundance and is more extensively r immed by covelli te with increasing elevation. Gold is found in t h e native s ta te . Silver occurs primarily in argentif erous covelli te but lesser amounts a r e also found in e i ther a rgen t i t e or acanthi te (not ye t determined), matildite, s t romeyer i te , and e lec t rum.

Late-stage mineralization a t Summitvil le i s character ized by a barite, jarosite, goethite, and gold assemblage found in t h e uppermost levels of t h e deposit. Sulfides associated with this assemblage a r e ext remely rare. A yellow phase included in a fluid inclusion in bar i te (Stoffregen, 1985) suggests t h e possibility t h a t nat ive sulfur may also be pa r t of th is assemblage. This late-stage assemblage, although volumetrically minor, locally contains high grades of gold.

Supergene oxidation extends t o a depth of 100 t o 200 f e e t below t h e surface. Copper is essentially complete ly removed from the oxidized zone, immedia te ly below which digenite and lesser chalcoci te c o a t and replace other sulfides.

Geochemical Environment

Defining the geochemical environment of t he Summitvil le deposits is l imited by t h e cu r ren t lack of fluid-inclusion and isotope data. Pa ramete r s such a s t empera tu re of mineral deposition, chemical composi- t ion of t h e fluids, paleodepth, and origins of t he fluids and dissolved consti tuents a r e f a r less well constrained than a t Creede. However, Stoffregen (1985) has done an excel lent job of limiting t h e conditions of a l tera t ion and mineralization using equilibrium thermodynamics. We will t ake t h e same approach, but we will also present some a l ternat ive interpretations.

Temperature - salinity ranges -- According t o Stof f regen (19851, t h e lack of fluid-inclusion d a t a is due t o a paucity of primarv fluid inclusions. He found . . only th ree samples with measurable primary inclusions in the small euhedral qua r t z crys ta ls intergrown with sulfides, lining voids in t h e vuggy sil ica and none were found in sphalerite o r alunite. His preliminary data , based on 19 measurements, show homogenization t e m e ra tu res in these samples ranged f rom 230' t o 8 320 C, with most of t h e values between 230' and 270 '~ . Two salinity measurements were 4 and 6 wt.-% NaCl equivalent. Perkins and Nieman (1983) also repor t t empera tu res fo r Summitvil le of 250' t o 280°c, presumably f rom fluid-inclusion measurements.

In contras t t o the lack of good primary inclusions, quar tz phenocrysts contain abundant secondary inclusions. The lack of secondary inclusions in quar tz phenocrysts outside of t h e vuggy silica and quartz-alunite a l tera t ion zones (Stoffregen, 1985) is consistent with Bruha and Noble's (1983) suggestion t h a t these inclusions represent t he fluids responsible for t h e intense quar tz + aluni te + pyr i te alteration. Bruha and Noble (1983) measured homogenization temperatures of 231' t o 276OC, and salinit ies of 7 t o 21 wt.-% NaCl equivalent (averaging 10 wt.-%) in these secondary inclusions f rom one sample. They also repor t a s many a s 5 (unidentified) daughter (or trapped) minerals present in some inclusions. Limited da ta col lec ted by G. H. Symmes in our laboratory a r e consistent with those of Bruha and Noble. Based on paragenet ic relations and on di f ferences in salinity between these secondary inclusions and t h e primary inclusions in the vuggy silica zone, Stoff regen (1985) suggests t h a t t he high-salinity fluids found in t h e f r ac tu res in quar tz phenocrysts represent the a l tera t ion fluids, but not t he l a t e r ore-forming fluids. These very l imited d a t a may indicate an evolution from high-salinity t o low-salinity wa te r s in t h e Summitvil le hydrothermal system, but many more measurements a r e needed t o substant ia te any such conclusion.

The high salinity of t h e fluids which preceded o re deposition is in agreement with observations made by Reynolds (Fluid, Inc., personal communication; also s e e Bodnar e t al., 1985, this volume). H e observed a few isolated, healed microfractures in some ear ly quar tz , defined by e i ther vapor-rich H 2 0 + C 0 2 inclusions and/or halite-bearing inclusions with smal l vapor bubbles. While these inclusions a r e thought by Reynolds to represent early magmat i c fluids, the i r genesis is st i l l uncertain due t o t h e l imited information.

Barite, which is an unreliable host fo r fluid- inclusion d a t a (Ulrich and Bodnar, 19841, i s t h e only hydrothermal mineral with relatively abundant fluid inclusions. Cunningham (1985) repor ts preliminary temperatures of roughly 1 0 0 ~ ~ fo r inclusions in this late-stage bar i te associated with the famous gold "boulder" found in talus slopes of t he South Mountain dome in 1975. No salinity d a t a were reported.

Paleodepth--Based on geologic reconstruction, Steven and R a t t e (1960a) e s t i m a t e t h e paleodepth t o the top of the o re between 150 and 400 meters . Using a t empera tu re of 250°C, a salinity of 10 wt.-%, and t h e lack of evidence for boiling f rom the l imi ted d a t a


on both secondary and primary fluid inclusions, a minimum depth of deposition of 400 me te r s is es t imated f rom t h e tables of Haas (1971).

Sources of constituents--No hydrogen-isotope d a t a a r e available for anv of t h e Acid-Sulfa te- tv~e

, a

deposits. Goldfield is t he o h y one with oxygen-isotope data , and therefore i t is used t o help docurnent t he source of fluids in this t pe of deposit. According to Taylor (19731, t he A1'O da ta from Goldfield a r e compatible with ore-bearing fluids of essentially 100% meteo r i c rigin. Taylor does note , however, t h a t t he overall depletion a t Goldfield is appreciably smaller than a t nearby Tonopah (an Adularia-Sericite deposit), and suggests t h a t a t Goldfield i t is likely t h a t e i ther t h e a l tera t ion occurred a t lower t em e ra tu res (125-20ooc), or from fluids with a higher 6'$ value. While Taylor favored the former , recent fluid- inclusion da ta from Bruha and Noble (1983) indicate temperatures of about 230' t o 270°C a t Goldfield, thus suggesting t h a t t h e fluids may have been richer in 1 8 0 implying a significant magmat i c component. In t h e absence of deuterium data , i t i s impossible t o do more than specula te on the source of t he Goldfield o re fluids. Well-constrained light-stable-isotope studies a r e badly needed on deposits of t h e Acid-Sulfate type.

Whitney (1984a,b) has ca lcula ted the sulfur speciation in quenched magmat ic gases evolved from rnagmas of various oxidation states. Relatively oxidized magmas such a s those a t Julcani (Drexler, 19821, and t h a t giving rise to t h e Fish Canyon ash-flow tuff in t h e Cen t ra l San Juan Mountains (Whitney and Stormer , 1983) produce gases rich in SO2. SO2 gas is unstable a t temperatures below 400°C in the presence of water (Iwasaki and Ozawa, 1960; Holland 1965,1967; Sakai and Matsubaya, 1977; Burnham, 1979) and disproportionates into sulfuric ac id and H2S gas. In our opinion, t h e intense acid-sulfate a l tera t ion character is t ic of this deposit t ype results from t h e the a t t a c k on t h e wallrock by the H2S04 produced by the disproportionation of SO2., Although Whitneyls calculations a r e consistent wlth a magmat i c source for t he sulfur, they do not demonstra te t h a t such a source was necessary.

Stoffregen (19851, taking a d i f ferent approach, has shown t h a t t he intense acid-sulfate a l tera t ion, which preceded ore deposition, can be produced by a fluid whose chemistry is dominated by magmat i c gases. He modeled the in teract ions between an ascending sulfur rich "maginatic" gas (idealized a s H20-C02-SO -H2S-HCI) and liquid water using equilibrium t<ermodynarnics. H e describes t h e model a s 'Ithe in teract ion of a large mass of the vapor phase with a relatively small amount of liquid, such t h a t t he chemistry of t h e liquid is completely controlled by t h a t of t he vapo.r." The results of his calculations show t h a t a t 2 5 0 ' ~ and vapor-saturation pressure, a solution equilibrating with a vapor phase consisting of subequal amounts of SO2 and H2S (totaling about 2 wt.-%) and minor HCI could produce t h e vuggy silica and acid-sulfate a l tera t ion seen a t Summitville.

As discussed in a previous section, both t h e lead and sulfur-isotopic da ta for Summitvil le and t h e other Acid-Sulfate deposits a r e similar isotopically t o t h e enclosing volcanic rock indicating t h a t lead and sulfur

were e i ther derived frorn the volcanics or from a re la ted magmat ic source.

Chemical parameters--Together with the above fluid-inclusion da ta , t h e a l tera t ion and ore mineralogy can be used t o put l imits on some of the chemical

Figure 7.18. Log aS2-%2 diagram at 2 5 0 ~ ~ showing the mineral stability fields for the significant minerals in the Summitville ores. Stof f regen's (1985) notation is used for the chalcopyrite-tennantite ore assemblage (IIa), and the covellite ore assemblage (IIb). An alternative to Stoffre- gen's interpretation for the chalcopyrite- bearing assemblage is shown as "IIa*". The boundaries for this alternative interpretation were calculated at log total sulfur = -3, while Stoffregen used a value of log total sulfur = -1. The heavy dashed lines show the alunite, kaolinite, and Kmica fields and the pH = 2 contour, all calculated at a log total sulfur concentration = -1; the heavy dotted lines show the kaolinite and Kmlca fields and the pH = 2 contour, all calculated at a log total sulfur concentration = -3 (the alunite field is not present at this low total sulfur concentration). The contour for 20 mole percent FeS in sphalerite coincides with the pyrrhotite field boundary. Quartz is stable throughout the diagram. The data for the iron-copper sulfides and oxides are summarized in Barton and Skinner (Table 7.2, 1979); the alunite-kaolinite line, interpolated from Hemley et al. (1969); the kaolinite, Kmica, Kfeldspar data are from Henley et al. (1984). Abbreviation: po = pyrrhotite.

paramete r s of t he hydrothermal fluid, such a s ac t iv i t ies of S2 and 0 pH, and to t a l sulfur concentra t ion during t h e diz ierent s tages of a l tera t ion and mineralization. As shown on Figure 7.18, ac t iv i ty of sulfur is probably t h e eas i e s t of these parameters to quantify. At Summitvil le, t h e quartz-alunite a l tera t ion and the l a t e r main-stage, upper-level, enargite-luzonite-covellite assemblage (s tage IIb) indicate an environment a t fairly high sulfur fugacities. The low silver concentrations (usually less than 1.0 wt.-%) which Stoff regen (1985) measured in e lec t rum from the unoxidized zone is consistent with such high sulfur fugaci t ies (Barton and Toulmin, 1964). The deeper assemblage of chalcopyrite and tennant i te (s tage IIa) specifies a n environment a t significantly lower sulfur fugacities. The act iv i ty of oxygen during t h e deposition of this deep assemblage is st i l l open to question. On Figure 7.18, field "IIa" represents Stoffregen's (1985) in terpre ta t ion and field "IIa*" represents a n a l ternat ive in terpre ta t ion which will be discussed la ter . In e i the r case, t h e sulfur fugacity was significantly lower during t h e deposition of the deeper chalcopyrite-tennantite assemblage than i t was during the upper, covelli te assemblage. The maximum iron content of 1.5% FeS in sphalerite l imits t he lower sulfur fugacity of t h e deeper assemblage t o values g rea te r than about -I I a t 250°C.

The calculations by Whitney (1985) show t h a t a n SO2-rich magmat i c gas quenched to 2 5 0 ' ~ would have a c t ~ v i t i e s of S and O2 in the alunite-stable field. Calculations wsich a r e consistent with those of Brimhall and Ghiorso (1983) and Stoffregen (1985), reinforce Stoffregen's conclusion t h a t t he fluids responsible for t h e acid-sulfate a l tera t ion a t Summitvil le resulted f rom t h e condensation of a magmat i t ic vapor plume in to deeply circulating meteor ic water.

Delineating the pH, total-sulfur concentration and oxygen act iv i ty is considerably more difficult. Figure 7.19 shows a ser ies of ac t iv i ty 02-pH diagrams a t d i f ferent total-sulfur concentrations relevant t o Summitville. Stoffregen's (1985) notation for t he d i f ferent mineral assemblages is used: vuggy silica (Ia), quartz-alunite-pyrite (Ib), low fS2 (chalcopyrite- bearing) o r e assemblage (IIa), and high fS2 (covelli te dominated) ore assemblage (IIb). As seen on Figure 7.19a, very acidic (pH <3.7) and oxidizing fluids a r e required t o produce the quartz-alunite alteration. Presumably, t he ex t r eme base leaching of t h e vuggy silica a l tera t ion was t h e result of even more acidic solutions. Stoff regen e s t ima tes t h a t a t a pH less than about 2 appreciable aluminum mobility would inhibit alunite deposition, result ing in a solution t h a t would dissolve most rock-forming minerals excep t quar tz and pyrite, c reat ing t h e observed vuggy silica assemblage. The a l tera t ion sequence of vuggy silica 4 quartz- alunite--quartz-kaolinite---illite + montmoril lonite is clearly t h e result of decreasing acidity away from the vein. Occasionally t h e quartz-kaolinite zone is absent and a transit ion from quartz-alunite t o i l l i t ic a l teration is observed. This may be t h e result of increasing salinity which causes t h e kaolinite field t o shrink (in the sul fa te portion of t h e diagram) and and finally disappear a t a potassium act iv i ty of approximately 0.1s a t a t o t a l sulfur concentration of 0.1 molal.

*Stoffregen (1985) ca lcula tes t h e disappearance of the kaolinite field a t a potassium act iv i ty of 0.005; the d i f ference between our calculations is due t o d i f ferent thermodynamic da ta (see Henley e t al., 1984, p. 81, for discussion of thermodynamic d a t a on the kaolinite- Kmica reaction). Assuming 0.005 a s t h e maximum potassium activity, Stoff regen uses a potassium act iv i ty of 0.0001 (log = -4) in his calculations. We feel t h a t is an unreasonably low value based on the few fluid-inclusion salinity measurements which have been done (ranging from about 1 to 3 molal) and based on our calculations for t h e disappearance of t h e kaolinite field. We have used a log potassium value of -1.5 based on a I molal solution with a Na/K concentration r a t io of 10.

Based on t h e r a re occurrence of nat ive sulfur associated with alunite, t h e total-sulfur concentration during the acid-sulfate a l tera t ion was probably between 0.1 and 0.01 molal. Figure 7.19b shows t h e re la t ive a f f e c t s on the sulfur and a luni te fields caused by reducing t h e total-sulfur concentration. Total- sulfur concentra t ions were probably never much g rea te r than 0.1 molal, because a t those concentrations the nat ive sulfur field expands greatly and would have been a dominant phase in t h e assemblage.

Compared to the deeper, s t age IIa assemblage, t h e covelli te-dominated s t age IIb assemblage is fairly res t r ic ted in t e rms of oxygen act iv i ty and pH (Fig. 7.19). The log total-sulfur concentration was probably between -1 and -2, similar t o the acid-sulfate a l tera t ion, based on the local occurrence of sulfur associated with these ores.

The geochemical environment during deposition of the chalcopyrite-tennantite assemblage, t he (deeper) s t age IIa ores, i s not a s well constrained a s t h a t of t h e a l tera t ion s tages (Ia and Ib) or t h e higher s t age IIa ores. As discussed above, there is no doubt t h a t this assemblage formed a t lower ac t iv i t ies of S2 than t h e covelli te-enargite assemblage. However, a s seen on Figure 7.19, t h e chalcopyrite-tennantite assemblage is s table over a wide range of oxygen activit ies. I t is bounded by t h e chalcopyrite = pyr i te + bornite react ion a t high S2 act iv i t ies and by the 1.5% FeS in sphaler i te contour a t low S2 activit ies. The possible environment for t he chalcopyrite-tennantite assemblage is fur ther res t r ic ted by Stoffregenls observation t h a t this assemblage (as well a s t he covelli te assemblage) is associated with minor kaolinite. There a r e two a reas on the a. pH diagram where these conditions a r e m e t (Figs. 519a -c ) ; t he field labeled "IIa", a t lower oxygen activit ies, corresponds t o Stoffregen's in terpre ta t ion and field "IIa*", a t higher oxygen activit ies, but lower to t a l sulfur, represents an a l ternat ive interpretation. Stoffregen (1985) chose the lower oxygen-activity field because t h e IIa* field borders on t h e hematite-stabil i ty field and no primary hemat i t e has been identified with t h e unoxidized ore. Even so, we feel t h a t th is upper field represents a likely environment because: ( I ) i t has values of oxygen act iv i ty and pH similar t o both the preceding acid-sulfate a l tera t ion s t age and the essentially cogenetic, but higher, s t age IIb ores, and


-29

-30 L o g Cs = - 3

-3 1

-32

-33 (V

0 -34 a n, 0 -35

-36

-37

Figure 7.19. A series of log a -pH diagrams constructed for 2 5 0 ~ ~ at dygferent total sulfur concentrat ions relevant to Summit- ville. The salinity is 1 molal, with Na+/K+ = 10. Stof fregen's (1985) notation for the different mineral assemblages is used: vuggy silica (Ia), quartz-alunite- pyrite (Ib), low f S2 (chalcopyrite-bearing) ore assemblage (IIa), and high f S2 (covellite dominated) ore assemblage (IIb). An alternative to Stof f regen's interpretation for the chalcopyrite-bearing assemblage is shown as "IIa*". Log total sulfur for A). = -1; for B). = -2; and for C). = - 3. See Figure 18 for sources of data. Abbreviations: ten = tennantite, eng = engarite, cov = covellite, dig = digenite, py = pyrite, bn = bornite, cpy = chalcopyrite, p = ~rrhotite.

(2) i t l ies in t h e sulfate-dominant portion of t h e t h e higher oxygen-activity field (IIa*). This can be diagram which is consistent with t h e introduction of a resolved by a model which calls on lower total-sulfur sulfur-rich magmat i c gas a s discussed above (Whitney, concentra t ions deeper in t h e deposit (discussed below). 1984a,b; Brimhall and Ghiorso, 1983). As can be seen in Figures 7.19b and c, t h e t ennan t i t e

One problemat ic a spec t of this in terpre ta t ion is f ield shi f t s t o lower pH's with decreas ing to t a l sulfur. format ion of t ennan t i t e associa ted with kaolinite in Two o the r f ac to r s which favor t ennan t i t e associa ted

with kaolinite are: (1) Stoff regen repor ted minor Sb substi tution in tennant i te which appears t o increase with elevation in t h e deposit , and (2) t h e possible trend of decreasing salinity (potassium activity) with t ime which was discussed earlier.

As seen in Figure 7.19, excep t for being slightly less acidic, t he chemistry of fluids which deposited the covelli te assemblage was very sirnilar t o those which produced the acid-sulfate a l tera t ion. Because the base of t h e vuggy silica zone roughly corresponds t o the base of t he covelli te assemblage (Fig. 7.171, i t is possible t h a t similar processes may have a f f ec t ed them both. During the a l t e ra t ion s tage , Stoffregen (1985) in terpre ts t he base of t h e vuggy sil ica zone a s t h e level of in teract ion between the magmat i c gases and water. If t he magmat i c gases continued t o be condensed a t this s ame level during o re deposition, i t is consistent t h a t t he higher sulfur assemblage begins here. Below this level (which presumably would vary with time), t h e tennant i te assemblage would form a s a consequence of the lower total-sulfur concentrations and activit ies. The di f ference between t h e alteration- and ore-deposition s tages may be r e l a t ed to the increase in permeabili ty following the acid-sulfate leaching which allowed fo r increased fluid flow resulting in a greater dilution of the magmat i c gas by rneteoric waters.

Summary of Summitvil le ~Vineralization - The work by Stoffregen (19851, Steven and ~ a t t ;

(1960a), Mehnert e t al. (1973) and Perkins and Nietnan (1983) has shown t h a t t h e Summitvil le ores were deposited shortly a f t e r t h e emplacement of t h e host volcanic dome on the margins of t he Summitvil le and Pla toro calderas. This temporal and spat ia l magmat ic association obviously a f f e c t e d the chemis t ry of t he hydrothermal system. Stoff regen (1985), Whitney (1984a,b), and Brimhall and Ghiorso (1983) have shown t h a t t h e intense acid-sulfate a l tera t ion, which preceded o re deposition, can be produced by a fluid whose chemistry is dominated by magmat i c gases. Stoffregen suggests t h a t t h e base of t h e vuggy silica a l tera t ion represents t h e level of in teract ion of these magrnatit ic gases. Ore deposition, which followed, occurred a s two di f ferent mineral assemblages, a deeper, chalcopyrite assemblage, and an upper, covelli te assemblage. The geochemical conditions which produced these o res is not well understood. Perkins and Nieman (1983) and Stoff regen (1985) have suggested t h a t t h e change from acid-sulfate leaching t o o re deposition ref lec ts a transit ion f rom a vapor- dominated geothermal system t o a liquid-dominated system. Stoffregen envisions the chemis t ry of the system t o then be dominated by me teo r i c fluids, which a r e less oxidized and less acidic. He also suggests t h a t t h e change in the sulfide assemblages with elevation is due t o t h e influx and mixing with acid-sulfate waters in to t h e upper portion of t h e deposit a s has been modeled by Reed and Spycher (1985, this volume). We ag ree t h a t t he change from al tera t ion t o o r e deposition represents decrease in the magmat i c vapor contribution to t h e system, but suggest t h a t such a vapor may sti l l have played a significant ro le in chemistry of t h e solutions depositing t h e ores. The change

f rom the lower t o t h e upper sulfide assernblage, which roughly corresponds to t h e base of t h e vuggy silica, may sti l l represent t he level of t he in teract ion of the magmat ic gases. With t h e present information, i t is not possible t o d i f ferent ia te which mechanism is more likely. The final s t age of mineralization, t he barite, jarosite, goethite, gold assernblage represents the waning s tages of hydrotherrnal ac t iv i ty and the collapse of the system.

GEOTHERMAL INTERPRETATION OF VOLCANIC-HOSTED EPITHERMAL DEPOSITS

Much of t h e conceptual base on which recent models of epi thermal mineralization have been built has come from observations on act ive geothermal sys tems (cf. White, 1981). Of particular importance have been observations on t h e anatomy of ac t ive sys tems and on t h e identification of processes operating within them. In th is synthesis we have a t t empted t o evaluate t h e observational base for epi thermal ore deposits in t e rms of t h a t for ac t ive systems. In summary, i t is useful t o combine observations on both ac t ive geothermal sys tems and well- studied ore deposits in a n a t t e m p t to in terpre t t he Adularia-Sericite and Acid-Sulf a te- type epi thermal deposits in a geothermal framework. In the f i rs t chapter in this volume, Henley has summarized most of t he pertinent observations on act ive systems; and the processes opera t ing within them have been explored in deta i l in the f i rs t volurne of this series (I-lenley e t al., 1984). The r e c e n t review by Henley and Ellis (1983) has established a geothermal in terpre ta t ion for hydrothermal ore deposits in general. The following discussion leans heavily on these sources and on the preceding discussion in this chapter without further docu~nenta t ion.

Adularia-Sericite Deposits

The observational base on Adularia-Sericite-type deposits suggests t h a t t hese deposits formed in hydrothermal systems similar t o those of the Taupo volcanic zone in New Zealand and t h e Valles and Yellowstone calderas in the United Sta tes , a s suggested by Henley and Ellis (19831, and discussed by Henley (1985, this volume). The general s t ruc tu re of such geothermal sys tems is i l lustrated in Figure l . l a in Henley's chapter (1985, this volume), which should be compared with the schemat i c representation of t h e Creede hydrothermal system a s i l lus t ra ted in Figure 7.11 of this chapter. From such a comparison t o modern geothermal systems, the following character is t ics of hydrothermal sys tems responsible for t h e format ion of Adularia-Sericite-type deposits appear to be particularly important: 1. The ore zone occurs a t distances of several

kilometers f rom t h e h e a t source t h a t drove t h e hydrothermal circulation. The ores occupy a narrow ver t ica l interval, t h e top of which lies a t distances of 200-600 m e t e r s below t h e paleo-water table. Au-rich deposits may have formed in a somewhat more shallow environment than did the Ag-rich deposits.


2. The ores appear t o have been deposited along the top of a deeply circulating, moderately t o strongly concentra ted brine a t , or close to , the in ter face of t h a t brine with overlying ground waters.

3. In the upper par t of t he brine plume, near i t s in ter face with t h e overlying ground waters, fluid flow appears t o have been dominated by l a t e ra l movement, and discharge t o t h e surface appears t o have been confined to small a r eas consti tuting only a small f rac t ion of t h e to t a l projected surface a rea of t he system.

4. Either lithologic aquitards or gravity-driven ground-water flow above t h e circulating brine plume have formed hydrologic caps on the plume forcing l a t e ra l flow across i t s top; overlying s t ructures a r e barren of mineralization and of ten show no evidence of vein filling or wallrock alteration.

5. The waters charging the deeply circulating brine were predominantly of me teo r i c origin, although some magmat ic contribution cannot be ruled out.

6. The fluids in t h e upwelling plume appear t o have been near neutra l in pH and t o have had thei r redox and sulfidation s t a t e s buffered by reactions with iron-bearing minera ls in t h e wallrock or in t h e vein assemblage.

7. Boiling occurs in the upper parts of t he brine plume, and mixing with surrounding ground waters occurs along the sides and, particularly, along the top of t h e plume. The depth a t which boiling commences is strongly dependent on the gas content of t he waters, and t h e degree of mixing dependent on t h e density contras ts between the plume and t h e surrounding ground waters. Both processes lead t o the precipitation of o re and gangue minerals.

8. The acid required t o produce the ser ic i t ic a l ter - ation associated with the deposits was produced by the condensation, in the upper par ts of the system, of acid volati les distilled off t h e deeply circulating brine.

9. The metals, and perhaps other components of the ores, appear t o have been derived by leaching of wall rocks deep in t h e systern.

Acid-Sulfate Deposits

In con t ra s t t o the fairly good database for t h e more numerous Adularia-Sericite-type deposits, t h e observational base for t h e Acid-Sulfate-type deposit is substantially more limited. There is also a similar, but less pronounced di f ference in the amount of da t a available on t h e two environments in ac t ive geothermal systems. However, we feel t ha t t h e d a t a bases on both t h e ore deposits and t h e ac t ive sys tems a r e sufficiently consistent t h a t we a r e comfor table in using t h e model of geothermal sys tems in andesi t ic volcanic t e r r anes presented a s Figure 1.lb in the chapter by Henley (1985, this volume) a s a basis for in terpre t ing Acid-Sulfate-type epi thermal ore deposits. 1. Observations on Acid-Sulfate-type deposits

indicate both a temporal and spat ia l genet ic association with the co re of a volcanic dome.

2. The ore zone occupies a narrow ver t ica l interval, t h e top of which appears t o have been within 200- 500 me te r s of t h e paleo-water table. There is some indication t h a t t h e Au-rich Acid-Sulfate-type deposits may have formed a t shallower levels than have the Ag-rich deposits.

3. Although t h e d a t a a r e quite l imited, both the sulfur- and lead-isotopic d a t a indicate t h a t they were e i ther derived directly from a magmat ic source or from t h e enclosing volcanic rocks.

4. Light-stable isotopic evidence on t h e origins of t h e fluids i s almost non-existent but can be in terpre ted t o provide evidence fo r a magmat ic contribution t o t h e fluids responsible for wallrock alteration.

5. The sulfuric ac id responsible for t h e intense acid- sul fa te wallrock a l tera t ion which preceded mineralization appears t o have been generated by the disproportionation of SO2 gas contained in the magmat i c vapor plume in to sulfuric ac id and H2S gas.

6. A surficial zone of acid-sulfate a l tera t ion due to t h e a tmospher ic oxidation of H2S t o form sulfuric ac id would also be expected in this type of hydrothermal sys tem, but has not yet been documented in any mining district .

7. The relatively high sulfidation s t a t e of t he fluids responsible for both the acid-sulfate a l tera t ion and t h e covelli te-enargite fac ies of t he sulfide mineralization i s a n expected consequence of the degassing of a relatively oxidized magma.

8. The increased permeabili ty produced by the intense base leaching associa ted with t h e acid-sulfate a l tera t ion provided channelways for subsequent ore-fluid migration.

9. The a l tera t ion and mineral assemblages of the Acid-Sulfate-type deposit, along with t h e significant copper contents and the association with porphyrit ic rocks, suggest t h a t this type of epi thermal deposit bears some genet ic relation t o porphyry-copper systems. The na tu re of such a relationship, if i t exists, remains speculative.

MECHANISMS OF ACID-SULFATE ALTERATION

Intense wallrock a l tera t ion t o an assemblage consisting of qua r t z + kaolinite + aluni te is t h e most pronounced character is t ic of Acid-Sulfate-type deposits. Most of th is a l tera t ion takes place in t h e o re zone a t depths of severa l hundred me te r s below the water table. However, t h e same assemblage is produced by sulfuric ac id a t t a c k on wallrocks in two very di f ferent surficial environments. I t is important t o have a s e t of cr i ter ia t h a t will allow us t o distinguish between these th ree di f ferent origins of acid-sulfate a l tera t ion so t h a t we may identify the environment in which any particular a l tera t ion zone formed. Cr i t e r i a based on pat tern recognition and/or textura l evidence can provide compelling evidence t o make such identification, but in most cases they do not and t h e r e has been considerable confusion in t h e l i t e r a tu re based on misidentification of the origin of acid-sulfate alteration. In most cases, co r rec t identification can be made with the application of light-stable isotope analyses, somet imes in conjunction

with K-Ar a g e determinations. In this section, we will discuss t h e use of isotopic cr i ter ia t o make such distinctions.

Bethke (1984) discussed the following th ree reactions, e a c h of which operates in a d i f ferent environment (as noted below in parentheses) and results in t h e format ion of sulfuric acid leading t o t h e format ion of acid-sulfate a l tera t ion

4S02 + 4H20 = 3H2S04 + H2S (pr imary hypogene) (6 )

H2S + 202 = H2S04 (p r imary supergene) ( 7 )

2FeS2 + M20 + 15/202 = 2Fe203'3H20 + 4H2SO4 ( 8 ) ( secondary super gene)

React ion ( 6 ) is, in our opinion, t h e most important mechanism by which sulfuric acid is formed in Acid-Sulfate-type o r e deposits. I t describes t h e disproportionation of magmatically exsolved SO2 gas in to sulfuric acid and H2S gas deep in t h e system, a s discussed in t h e preceding section. We have noted this deep environment of acid-sulfate a l tera t ion a s "primary hypogene" in recognition of the f a c t t h a t i t occurs deep within t h e hydrothermal system. A number of ac t ive geothermal systems associated with andesi t ic volcanism, particularly in island-arc environments, contain deep, high-temperature acid waters. The acidity and sul fa te content of these waters appear t o have been generated by such a disproportionation reaction deep in the system. These include t h e Tahuangtsui thermal area , Tatun volcanic region, Taiwan, (Chen, 19701, t he Matsukawa geothermal area , Japan, (Nakamura e t al., 1970), and several o ther geothermal a r e a s in Japan (Kiyosu and Kurasawa, 1983, 1984). A t Matsukawa, in tense acid- sul fa te a l tera t ion also occurs a t t he surface , but this near-surface a l tera t ion probably formed from the oxidation of hydrogen sulfide in t h e s team-heated waters ( react ion 7) a s discussed below.

React ion (7) i s t h e simple oxidation of hydrogen sulfide t h a t produces the spectacular sol fa tar ic a l tera t ion seen a t t he surface in ac t ive geothermal systems. Because this environment, although surficial , is an in tegra l pa r t of t he hydrothermal system, and because t h e a l tera t ion results from wall rock in teract ion with descending acid waters, we have noted i t a s "primary supergene." Day and Allen (1925) originally proposed this mechanism t o explain a l tera t ion seen in the Lassen area , and White (1957) discussed i t more extensively (see also Henley and Ellis, 1983; Henley, 1985, this volume; and Reed and Spycher, 1985, this volume). Oxidation of t h e H2S occurs when t h e vapor phase generated by boiling of t h e deep waters con tac t s t h e a tmosphere just above t h e wa te r table. The sulfuric acid so generated percola tes back into, and acidifies, t h e s team-heated ground waters which overly the deeply circulating hydrothermal cell. These surficial, acid, steam-heated waters a r e essential e lements of almost a l l high- t empera tu re geothermal systems, from which both Acid-Sulfate and Adularia-Sericite types of epi thermal o re deposits form. The a l tera t ion produced by th is mechanism is very intense, but is a surficial phenomenon, rarely extending down t o depths of even 50

me te r s below t h e surface. These surficial a l tera t ion zones a r e probably not preserved in many o r e deposits because they a r e very sof t , shallow, and easily eroded.

The l a s t react ion (8) involves the production of sulfuric acid during supergene a l tera t ion of sulfide ore. Because t h e oxidation an teda te s t h e hydrothermal sys tems (sometimes by millions of years) we have labeled th is environment "secondary supergene." Such oxidation has produced a luni te a t Creede, Goldfield and, undoubtedly, in many other districts. Because this t ype of acid-sulfate a l tera t ion has l i t t l e t o do with the format ion of t h e deposit i t is important t o be able t o distinguish i t f rom the former two mechanisms.

Sulfur isotopes provide a means t o d i f ferent ia te among t h e t h r e e environments outlined above. Field (1966) was t h e f i rs t t o use sulfur isotopes t o distinguish between hypogene and secondary alunite. Jensen et al. (1971) showed t h a t both hypogene and secondary a luni te occurred a t Goldfield, and t h a t t h e t w o had widely di f ferent sulfur-isotopic compositions. The a luni te judged t o be hypogene on field petrographic grounds was strongly enr iched in 38; compared t o t h a t formed by secondary processes (Fig. 7.20). The secondary alunite a t Goldfield (and many o the r districts) has t h e same range of sulfur-isotopic composition a s do t h e primary sulfides because, due t o k inet ic ef fects , t he re is l i t t le or n o sulf ur-isotopic f rac t ionat ion during low-temperature oxidation of sulfides. In contras t , hypogene a luni te formed a t t empera tu res above approximately 250°C is much heavier in sulfur-isotopic composition than coexisting sulfides because of t he rapid isotopic exchange and s t rong f rac t ionat ion between aqueous sulfide and sul fa te species a t moderate t o high temperatures , particularly in ac id environments (Ohmoto and Lasaga,

JULCANI 1

BARITE D l SULFIDES

-10 -5 0 +5 +10 +15 +20 +25

SUMMITVILLE ALUNITE

[24 SULFUR

SULFIDES I

@ , , , , , , , , , , , , , , , 3 , , , , , , , , , , , , , ~ -10 - 5 0 +5 +10 +15 +20 +25

4 QOLDFIELD

J [&ql SECONDARY ALUNITE

ALUNIE 2

1

0 -10 -5 0 + 5 + l o +15 +20 +25

DELTA "48

F i g u r e 7.20. 6 3 4 ~ h i s t o g r a m s f o r J u l c a n i , S u m m i t v i l l e , and Gold f i e ld . Data sources : J u l c a n i , Goode l l (1977); S u m m i t v i l l e , R 0. R y e (U.S.G.S., p e r s o n a l c o m m u n i c a t i o n , 1 9 8 5 ) ; G o l d f i e l d , J e n s e n e t al. (1971) .


I EXPLANATION Kaolinite

0 Halloysite I a Meteoric water 1

EXPLANATION Quaternary

Thermal water

Akatani

- 100

I HYDROTHERMAL / I

F i g u r e 7.21. a ) . 6D - 6180 d i a g r a m f o r k a o - l i n i t e and h a l l o y s i t e formed f rom supe rgene a l t e r a t i o n o f some J a p a n e s e orebodies, adapted from Marumo et al. (1982). Dashed l i n e s c o n n e c t m i n e r a l a n a l y s e s w i t h their c o f g e s ~ n d i n g s u r f a c e w a t e r s . b). 6D - 6 0 d l a g r a m f o r h y d r o t h e r m a l k a o l i n i t e s a n d f o r thermal w a t e r s n e a r the m i n e r a l local i t ies , a d a p t e d f r o m Elarumo et al . ( l q 8 2 ) . T h e d a s h e d l i n e s h o w s the 6 D - 6 0 r e l a t i o n o f k a o l i n i t e i n e q u i l i b r i u m

w i t h m e t e o r i c w a t e r a t 100%.

1982). Such a relationship has been documented a t Summitvil le and Julcani a s well a s a t Goldfield (Fig. 7.20).

Although sulfur isotopes can be used t o distinguish secondary alunite (reaction 8) f rom hypogene alunite formed by disproportionation of SO2 (react ion 61, they cannot help in d i f ferent ia t ing between supergene alunite formed through t h e oxidation of primary sulfides by weathering processes ( react ion 8) and t h a t formed by the oxidation of H2S a t

t h e water table over s team-heated wa te r s (reaction 7). When H2S is oxidized in the s team-heated waters, i t undergoes l i t t le sulfur-isotopic exchange with the sul fa te produced because of slow react ion kinetics a t relatively low temperatures (<lOoOc) , a s has been demonstra ted by Steiner and Raf t e r (1966) for Wairakei, and by Schoen and Rye (1970) for Yellowstone. Therefore, t he alunites fo rmed in both supergene and steam-heated water environments will both have sulfur-isotopic compositions approximately t h e same a s those of thei r coexisting sulfides.

In order to distinguish the alunites formed from steam-heated waters from t h e supergene alunite, i t is necessary t o look a t t he associated kaolinites. Taylor (1974) showed t h a t hydrogen and oxygen isotopes can be used t o d i f ferent ia te between supergene and primary kaolinite. Figure 7.21a shows a group of kaolinites and halloysites formed by supergene a l tera t ion of some Japanese orebodies (Marumo et al., 1982). These da ta plot very close t o t h e "kaolinite line" (Savin and Epstein, 1970; Lawrence and Taylor, 1971) which represents the locus of hydrogen and oxygen isotopic compositions for kaolinites formed by weathering and assumed t o be in equilibrium with meteor ic waters along the me teo r i c water line. Figure 7.21b shows the hydrogen- and oxygen-isotopic compositions of a group of kaolinites from acid-sulfate a l tera t ion zones from several geothermal a reas (solid circles) and o r e deposits (open circles) in Japan. Plots of t he oxygen- and hydrogen-isotopic compositions of these kaolinites, formed in t h e steam-heated water environment, l ie considerably removed from the meteor ic water line and very close to t h e dashed line which marks the compositions of kaolinites t h a t would be in equilibrium with me teo r i c waters a t temperatures of 1 0 0 ~ ~ .

Table 7.6 summarizes t h e isotopic discriminators current ly available for determining t h e origin of acid- sul fa te alteration. I t indicates t h a t sulfur isotopes can be used t o identify alunite formed by t h e dispropor- t ionation of SO2? and t h a t t h e hydrogen- and oxygen- isotopic cornposltion of kaolinites can be used t o distinguish between primary and secondary supergene origins for shallow acid-sulfate a l tera t ion zones. In addition, K/Ar a g e determinat ions on a luni te can also be used in many dis t r ic ts t o distinguish between supergene and primary origins, a s has been done a t Creede (M. Lanphere, personal communication, 1981) and Round Mountain (Tingley and Berger, 1985).

Recent advances by Pickthorn and O'Neil (1985) and by Bethke, Rye, 'Wasserman, and Goss (P. Bethke, personal communication, 1985) on the se lec t ive analysis of t he isotopic composition of oxygen in both the sul fa te and hydroxyl s i tes in a luni te is expected t o allow t h e distinction between alunites formed in t h e s team-heated versus the supergene environment t o be made on a luni te i tself , obviating t h e need to use associated kaolinite. Such application, however, awai ts fur ther definition of t h e sys temat ics of oxygen- and hydrogen-isotopic f rac t ionat ion between alunite and water , and t h e demonstration of i t s feasibility.

I t should be noted t h a t in Acid-Sulfate deposits such a s Summitvil le we may expec t t o encounter acid- sul fa te a l tera t ion formed by a l l t h r e e mechanisms discussed above. Although not ye t documented for any

CHAPTER 7

T a b l e 7 .6 - -Cha rac t e r i s t i c s of a c i d - s u l f a t e a l t e r a t i o n o f d i f f e r e n t o r i g i n s (Bethke , 1984)

O x i d a t i o n D i s p r o p o r t i o n a t i o n Supergene of H2S o f SO2 o x i d a t i o n

-- -

Reac t ion ( T a b l e 4 ) 1 2 3

3 4 ~ a l u n i t e = s u l f i d e s s u l f i d e s = s u l f i d e s

180 k a o l i n i t e removed from removed from n e a r k a o l i n i t e l i n e k a o l i n i t e l i n e k a o l i n i t e l i n e

K/Ar age a l u n i t e concordan t w i t h concordan t w i t h d i s c o r d a n t w i t h m i n e r a l i z a t i o n m i n e r a l i z a t i o n m i n e r a l i z a t i o n

d is t r ic t , we should e x p e c t t h e format ion of primary- supergene acid-sulfate a l t e r a t ion (reaction 7) in t h e shallow pa r t s of t h e hydrothermal system, probably sepa ra t ed f rom t h e deepe r hypogene acid-sulfate a l te ra t ion . In t h e f e w dis t r ic ts where sulfur-isotope d a t a on a luni te and associa ted sulfides a r e available most of t h e a luni te appea r s t o have been formed e i ther by react ion (6) in t h e d e e p a l tera t ion zone or f rom supergene oxidation of pr imary sulfides ( react ion 8). However, Henley (1985, t h i s volume) points ou t t h a t lakes severa l hundred m e t e r s deep filled with ac id sulfate-chloride wa te r may form in t h e c r a t e r s of s t ra tovolcanoes in andes i t ic volcanic terranes. Henley c i t e s a s a n example L a k e Ruapehu in New Zealand which is g rea t e r t h a n 300 m e t e r s deep and conta ins a sulfate-chloride wa te r of pH = 1.25 a t 55OC. The acidity and chemical composit ion of these lakes is due t o t h e f a c t t h a t t hey a r e f e d by volcanic vapors f rom below a s well a s by me teo r i c wa te r s a t t h e surface. In t h e ca se of such deep, acid, sulfate-bearing lakes, downward percolation of t h e lake waters through t h e underlying volcanic pile would produce in tense acid- su l f a t e a l t e r a t ion a t considerable depths beneath t h e lake (and summit) level. Such downward percolation might be expec ted t o occu r a s t h e deep hydrothermal sys tem begins t o wane, and i t is possible t h a t a n ear l ie r deep acid-sulfate a l t e r a t ion zone might be overprinted by a l a t e r , descending, acid-sulfate alteration! However, we cont inue t o believe t h a t t h e bulk of t h e acid-sulfate a l t e r a t ion in Acid-Sulfate-type o r e deposits resul ted f rom t h e format ion of sulfuric acid by t h e disproportionation of magmat i c SO gas.

We should a lso no te t h a t reac t ions (7) a n 3 (8) can a lso occur in Adularia-Sericite systems. Specifically, because surficial zones of acid-sulfate a l t e r a t ion due t o t h e oxidation of H2S in t h e s team-heated water zone a r e so common In ac t ive geothermal sys tems thought t o be analogous t o fossil sys tems represented by Adularia-Sericite-type epi thermal o r e deposits t h a t we should expec t t o s e e t h e m represented in Adularia- Sericite-type epi thermal distr icts. However, t o our knowledge no such occurrences have y e t been documented in any such dis t r ic t ; perhaps because they have been removed by erosion, perhaps because they have no t been specifically sought. However, secondary

a luni te formed a s a resul t of t h e supergene oxidation of pr imary sulfides has been documented by both K/Ar a g e determinat ions and sulf ur-isotope analyses a t both Round Mountain and Creede.

ACKNOWLEDGMENTS

We a r e gra teful fo r t h e numerous thought- provoking discussions on epi thermal and ac t ive geothermal sys t ems with Paul Barton and Dick Henley. We a lso benef i t ted from ta lks with Roger Stof f regen, whose comprehensive study a t Summitvil le, Colorado was essential fo r using Summitvil le a s t h e archetype for Acid-Sulfate epi thermal systems. Work done by J im Goss on t h e original compilation of geological d a t a fo r t h e sixteen epi thermal d is t r ic ts studied in de ta i l is gratefully acknowledged. In addit ion, we t a k e t h e full weighty responsibility f o r our s t a t e m e n t s and in terpre ta t ions , and we relieve our colleagues of any such burden.

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Chapter 8 GEOLOGIC CHARACTERISTICS OF SEDIMENT-HOSTED, DISSEMINATED

PRECIOUS-METAL DEPOSITS IN THE WESTERN UNITED STATES William C . Bagby and Byron R. Berger

INTRODUCTION

Sediment-hosted precious-metal deposits a r e typically formed in carbonaceous, silty dolomites and l imestones or calcareous sil tstones and claystones. Cold mineralization is disseminated in the host sedimentary rocks and is exceedingly fine grained, usually less than one micron in s ize in unoxidized ore. Pr imary a l tera t ion types include silicification, decalcification, argillization, and carbonization. Supergene a l tera t ion is dominated by oxidation resulting in the formation of numerous oxides and sul fa tes and t h e re lease of gold f rom i t s association with sulfides. Commonly associated t r a c e e l emen t s a r e arsenic, barium, mercury, antimony, and thallium. Deposits of this type a r e commonly referred t o a s e i the r Carlin-type deposits, a f t e r t h e large bulk- minable, disseminated-gold deposit in northern Nevada, or a s fine-grained or llinvisible-goldll deposits. We refer t o deposits of this t ype a s sediment-hosted, disseminated precious-metal deposits.

This chapter presents a classification scheme and reviews the geologic character is t ics of sediment- hosted, precious-metal deposits. The influences of geology on both mining and t h e development of gene t i c and exploration models a r e discussed. Although deposits of this type occur throughout t h e western United Sta tes , t he larges t concentra t ion of deposits and also t h e best understood a r e in Nevada. We have chosen, therefore , t o use se lec ted individual deposits f rom Nevada a s type examples t o support t h e classification scheme and t o provide t h e s tudent with an understanding of t h e similarit ies and di f ferences t h a t occur in these deposits. This chapter i s thus designed t o develop and nur ture t h e knowledge of t h e comparat ive geology of sediment-hosted, disseminated precious-metal deposits. This is accomplished by reviewing and comparing regional-, district- , and deposit-scale geologic characterist ics.

CLASSIFICATION

Deposits of this type display a var ie ty of geologic and geochemical character is t ics t h a t may be used t o divide t h e complete s e t of deposits in to subsets (Tables 8.1 and 8.2). The Carlin deposit has lent i t s n a m e t o this deposit t ype since i t s discovery in t h e ear ly 1960's. Early investigators realized t h a t t h e Carlin deposit was possibly typical of a c lass of deposits t h a t had been mined in t h e past , but for which

no classification scheme seemed particularly appropriate. These older known deposits a r e Cetchell , Nevada (Joralemon, 19511, Mercur, Utah (Butler e t al., 1920), and Gold Acres, Nevada (Cilluly and Masursky, 1965). As more of t hese deposits were discovered and developed in t h e 19701s, i t became apparent t ha t Carlin was possibly a n end member of t h e sediment- hosted, disseminated precious-metal deposit type. Thus, "Carlin-type deposit" is a misnomer for some sediment-hosted, disseminated deposits. For example, Carlin is hosted by sil ty dolomites whereas other deposits a r e hosted in shales and siltstones. In addition, t he gold o r e a t Carlin occurs in pod-like zones in favorable host lithologies adjacent t o ver t ica l faul ts and i t is commonly difficult t o visually distinguish o re from unaltered host rock. On t h e other hand, t he o re a t some other deposits is more easily distinguished from unaltered rock because of d i rec t association with e i ther in tense sil icification in the form of "jasperoids" (e.g., Pinson) or with a noticeable increase in silica veining (e.g., Preble). Although "Carlin-type deposits" a r e considered t o have ext remely high gold t o si lver ratios, ce r t a in deposits included in th is type actual ly have high silver values, with gold a s the major me ta l (e.g., Dee); o thers have high silver values and lack gold (e.g., Taylor), but re ta in similar associated t r a c e e l emen t s (Sb, As, Hg) and a l tera t ion types.

On t h e basis of t h e above character is t ics , we have defined t w o deposit-type subsets: t h e jasperoidal- and Carlin-type subsets (see Tables 8.1 and 8.2) of which t h e r e a r e gold-rich and silver-rich end members. Jasperoidal-type deposits a r e those wherein t h e majority of t h e gold and/or silver i s hosted in jasperoid* or in qua r t z veins and r e l a t ed silicified wall rocks. On t h e other hand, Carlin-type deposits a r e those wherein t h e gold and/or si lver appears t o be evenly distributed in t h e host rocks which do not always appear t o be silicifed. Ore zones in Carlin-type deposits a r e commonly pod-like and ex tend up t o tens of me te r s away f rom fau l t s whereas o r e zones in t h e jasperoidal-type deposits a r e most commonly limited t o narrow, shear zones. There is complete gradation between these two subset types and a deposit classi- f ied a s jasperoidal may have exploration potential for Carlin-type extensions and vice versa. This classification serves a s a means of examining and understanding the differences t h a t occur between and within t h e individual sediment-hosted, disseminated precious- me ta l deposits. In addition, t h e separa t ion is useful a t the deposit sca le for understanding genesis and is a useful concept in regional and dis t r ic t exploration programs.

Table 8.1--Geological aspects of selected sediment-hosted, disseminated precious-metal deposits in the Western United States.

Host rocks Igneous rocks

Composition and Deposit Formation Age Lithology Age occurrence

Alligator Pilot Shale Devonian- Thin bedded, Tertiary Siliceous, pumiceous Ridge Mississippian calcareous, tuff and younger

carbonaceous basaltic andesite siltstones lava flows and clay stones

Carlin Roberts Mountains Formation

Cortez Roberts Mountains Format ion

Dee

Silurian to Laminated, silty Cretaceous Altered quartz diorite Early Devonian to sandy, 130 m.y. and diorite dikes

carbonaceous, dolomitic limestone

Silurian to Laminated, silty, Oligocene Altered dikes and sills Early Devonian argillaceous, 33-35 m.y. of quartz latite; Mill

carbonaceous, Jurassic Canyon granodiorite pyrite-bearing 150 m.y. stock limestone with dolomite

Devonian Devonian Massive, Limestone and Ordovician fossil-rich Vinini Formation limes tone

Getchell Preble Formation Cambrian Phyllitic Comus Formation Ordovician shale with

interbedded limes tone

Mesozoic(?) Altered dikes in the Tertiary(?) Dee mine; intermediate

composition

Cretaceous Osgood Mountains pluton; Granodioritic with associated intermediate porphyritic dikes. All are altered and in part mineralized.

Gold Acres Roberts Mountains Silurian to Carbonates, argil- Cretaceous Altered and mineralized Format ion Early Devonian lites, and siltstones: dikes of intermediate Valmy Formation Ordovician All mineralized rocks composition; Wenban Formation Devonian occur as fault blocks ~ertiaryc?) @art2 latite sills

low in the upper plate that are altered. of the Roberts Mountains thrust.

Horse Canyon Vinini Formation Ordovician Siltstones and chert; Tertiary Altered dikes and sills Wenban Formation Devonian silty carbonaceous of quartz latite

limestone

Jerritt Hanson Creek Ordovician Carbonaceous, shaly Mid-Tertiary(?) Small dikes and plugs of Canyon limestones with chert diorite and a small

dolomite, and rhyodacite flow 2.4 km SW bioclastic limestones and 3 km NE of mineral-

ization, respectively

W. C. BAGBY & B. R. BERGER

Structure

Faults Folds Mineralization age Reference

N-NE, N-NW, and E-W trending N-NE striking No direct date of Klessig, 1984. normal faults. The NE asymetrical anti- mineralization. Ilchick, 1984. trending Vantage fault cuts form plunging to Tertiary tuff may be Ainsworth and Tertiary tuff. All faults SW. Extensive altered by gold system Brimhall, 1983. post antiform. fracturing and placing an lower age

brecciation along constraint. crest. Age unknown.

Devonian-Mississippian thrust NW directed folds fault (~oberts Mountains thrust), (antiforms and high angle faults trending E, N, synforms). Major and NE and NW. Faults are NW-trending pre-mineral with some post- antiform in the mineral movement. district. Age of

folding is Mesozoic.

No direct date of Adkins and Pota, mineralization. Post- 1984. mineral rhyolite lavas Hausen, 1967. and domes dated at 14 Radtke, 1981. m. y. provide upper age constraint and altered Cretaceous dikes (130 my.) provide lower age constraint.

N, NW, and EW trending, high Drag folds associa- Altered and mineralized Wells, Elliott, and angle, normal faults. Roberts ted with faults 33-35 m.y. old quartz Obradovich, 1971. Mountains thrust surrounds sympathetic to latite dikes place a Wells, Stoiser, and district. Roberts Mountains lower age constraint. Elliott, 1969.

thrust. NW direc- William C. Bagby, U.S. ted regional folds. Geological Survey,

unpub. field notes.

N, NW, NE, and EW trending high angle faults of ~esozoic(?) and Tertiary age. Roberts Mountains thrust (Dev.-Miss.).

N. trending Getchell fault zone includes several strands. Inception of fault Late Cretaceous.

N, NE faults dip steeply west. Roberts Mountains thrust.

NW directed folds Unknown, altered Wallace and Bergwall, of presumed dikes remain undated 1984. Mesozoic age. William C. Bagby, U.S.

Geological Survey, unpub. field notes.

Fold axis plunges Sericite in mineral- Joralemon, 1951. 45' NE. On south- ized granodiorite Silberman, Berger, ern limb sediments dated by K-Ar between and Koski, 1974. strike N and 87-92 m.y. This is Berger and Taylor, dip SE. inferred age of gold 1980.

mineralization.

NW directed anti- Not directly dated. Wrucke, 1984. forms and synforms Altered Tertiary(?) Wrucke and of presumed sill places a lower Armbrustmacher, 1975. Mesozoic age. age constraint.

N, NW, EW trending high angle NW directed folds. Altered Tertiary Coppinger and normal faults of Tertiary age. (Oligocene?) dikes Cartwright, 1983. Roberts Mountains thrust. place a lower age William C. Bagby, U.S.

constraint on Geological Survey, mineralization unpub. field notes.

Normal faults striking E-W, N-S Regional E-W Unknown N. 20-30 E. Roberts Mountains trending folds. thrust fault. NS faults post- mineral and high angle.

Hawkins, 1973; 1982. Stevens and Hawkins, 1984.

172 CHAPTER 8

Table 8.1--Geological aspects of selected sediment-hosted, disseminated precious-metal deposits in the Western United States--(continued).

Host rocks Igneous rocks

Deposit Format ion Age Lithology Composition and

Age occurrence

Mercur Lower Great Blue

Mississippian Massive, bedded Tertiary Fine-grained, porphyritic limestone; local 31.6 m.y. rhyolite plug south of bioclastic micrites deposit. Believed to and wackestones with post-date gold mineral- sparse siltstones ization. Also coarse-

grained porphyritic plugs 1.6 km north of deposit.

Northumberland Roberts Silurian Laminated silty Jurassic Altered tonalite and Mountains limestones, shales, granodiorite dikes and Format ion and siltstone pluton

Tertiary(?) Unaltered siliceous tuffs and altered rhyolitic dikes

Preble

Rain

Comus Formation Cambrian- Thin-bedded Cretaceous Intermediate composition Ordovician limestone and (90 m.y.) dikes and Osgood

shale Mountains pluton

Preble Formation Cambrian Phyllitic shale Cretaceous(?) Altered dike of and interbedded intermediate limestone composition

Webb Formation Mississippian Siltstone, shales, Tertiary and fine-grained (35 my.) sands tone

Sterling Wood Canyon and Cambrian Silty and sandy Tertiary Bonanza King dolomite with minor Formations carbonaceous matter

Taylor Guilmette Devonian Limestone and shaly Tertiary Format ion limestone (35 m.y.)

Quartz monzonite stock 6 mi south of deposit

Quartz latite dikes occur near the deposit.

Rhyolitic dikes altered but not mineralized

Tolman Pennsylvanian Sandy and calcareous Tertiary Altered intermediate to siltstones, silty mafic dikes and remnants limestone, and or rhyolitic ash-flow claystone tuffs. Rhyolitic domes

present 5 km SW.

*The t e r m jasperoid and jasperoidal a r e used repeatedly in t h e t e x t a s a rock name and a n adject ive , respectively. We use t h e t e rm a s an in tegra l pa r t of our classification of deposits of th is t ype and have adopted t h e definition of Lovering (19721, p. 3): "(1) jasperoids a r e composed predominantly of silica, which in most places is in the form of aphani t ic t o fine- grained quar tz , and (2) jasperoids form by replacement of t he enclosing rock."

REGIONAL GEOLOGIC CHARACTERISTICS OF DEPOSITS IN MINERAL TRENDS AND

ISOLATED DEPOSITS

Locations of major known sediment-hosted, disseminated precious-metal deposits in the western United S ta t e s a r e given in Figure 8.1. Except for Mercur, Utah, t he mos t thoroughly studied sedirnent- hosted precious-metal deposits a r e in Nevada; therefore , we have chosen severa l of t hese deposits for discussion in this paper. Some of t h e Nevada deposits occur along recognized mineral bel ts t h a t contain

W. C . BAGBY & B. R. BERGER

Structure

Faults Folds Mineralization age Reference

Northwest trending Late Cretaceous to late Oligocene Ophir anticline and Pole Canyon syn- cline. Thrust faulting was associated with folding.

Unknown. Likely Tertiary as normal faults are mineralized. Porphyritic plugs are unmineralized.

Tafuri, 1976. Kornze and others, 1984. Edwin W. Tooker, U.S. Geological Survey, oral commun., 1984.

(1) Normal and strike slip faults associated with folding. (2) Normal faults due to Basin and Range extension, (3) Normal faults of minor displacement possibly associated with intrusions.

Thrust faults (Dev.-Miss.), NE, N, to NW trending high angle Tertiary and Late Cretaceous faults.

Doming of Paleozoic sediments near Mesozoic pluton and folding of sediments near thrust faults

Altered dikes emplaced between 32-26 m.y. which places lower age constraint provided alteration due to gold hydrothermal system.

Ott, 1983. McKee, 1974. William C. Bagby, U.S.


NE trending high angle fault of Cretaceous(?) age

NE plunging ant iform.

Not dated but presumed to be the same as Getchell (87-92 my.).

Kretschmer, 1984. Berger, 1980 William C. Bagby, U.S.


Kretschmer, 1984. Berger, 1980. William C. Bagby, U.S.


N and NE to EW trending high angle faults. Cretaceous(?).

NE directed antiforms and synforms.

Not dated but presumed to be the same as Getchell (87-92 m.y.).

N and WNW trending high angle faults.

N-NW plunging antiform.

Unknown. Knutsen and West, 1984. William C. Bagby, U.S. Geological Survey, unpub. field notes.

Thrust fault contact between Wood Canyon and Bonanza King Formations.

Unknown. William C. Bagby, U.S. Geological Survey, unpub. field notes.

N, NW, NE trending high angle faults of Tertiary age.

Possible S plunging antiform.

The age constraints are provided by post mineralization, mid- Tertiary dikes, and early Tertiary, pre- mineralization faults.

Drewes, 1967. Lovering and Heyl, 1974. Havenstrite, 1983.

Low angle and high angle normal faults. No preferred direction.

Large and small scale folds trending N and NE.

Unknown. Brady, 1984.

several deposits of th is type whereas o thers a r e apparently isolated, individual deposits t h a t have not yet been demonstra ted t o occur with o the r similar deposits (Fig. 8.2). The geology of these bel ts and of the a reas around isolated deposits provides an understanding of the regional geologic character is t ics of these deposits.

The most famous gold bel t in Nevada is t he Carlin trend. This mineral bel t was initially referred t o a s the Lynn-Railroad belt and was described by Rober ts (1966). This bel t includes the Rain, Gold Quarry, Maggie Creek, Carlin, Bullion Monarch, Blue

Star , Gold Strike, Bootstrap, and Dee sediment-hosted, disseminated precious-metal deposits. Some investigators extend t h e Carlin bel t t o the southeast t o include the Alligator Ridge deposit (Fig. 8.2). However, detailed s tudies of t h e Alligator Ridge a rea and i t s relation t o extensions of s t ruc tu res within the Carlin trend a r e needed t o determine whether or not t he re is an association. The Horse Canyon, Cor tez , Cold Acres, and Tonkin Springs deposits occur along the Bat t le Mountain-Eureka bel t (Roberts, 1966) which is now more commonly referred t o a s t h e Cor t ez trend. The Getchell , Pinson, and Preble deposits

Table 8.2--Characteristics and form of ore in selected sediment-hosted, disseminated precious-metal deposits in the Western United States.

Alteration Ore bodies

Deposit Hypogene Supergene Form

Alligator Ridge

Carlin

Cortez

Dee

Getchell

Decarbonization, Oxidation decalicification, silicification, carbon remobilization, acid oxidation( ? )

Pods localized near high angle faults and extending into sediments

Decarbonization, Oxidation, calcite Pods localized along high angle faults silicification, remobilization, clay and extending into sediments calcification, carbon format ion remobilization, acid leaching with oxidation(?)

Silicification, acid Deep oxidation resulting Elongated zones paralleling leaching and oxida- in some redistribution faults and dikes and notably tion(?), decalcifi- of gold localized in breccia zones cation, and dedolomiti- associated with folds zation

Silicification and Oxidat ion argillization

Decarbonatization with Minor oxidation silicification and argillization; early calc-silicate skarnification

Gold Acres Silicification, early Deep oxidation contact metamorphism

Zones localized along faults

Sheet-like zones localized along strands of Getchell fault and pods in fold hinges

Tabular, dipping to SW parallel to thrust faults

Horse Canyon

Decalcification, Shallow oxidation Localized zones along NNE fractures silicification, carbon and in permeable hosts near mobilization fractures

Ore bodies

Gold or Mineralogy silver site Veins Reference

Oxidized: metallic gold, specular hematite, jarosite, stibiconite, goethite, quartz, barite, calcite, gypsum, alunite, and kaolinite. Unoxidized: stibnite, pyrite, orpiment, realgar, and calcite.

Oxidized: metallic gold, goethite, illite, kaolinite, barite, anhydrite, alunite, dolomite, calcite, quaartz, schuetteite, cinnabar, arsenolite, scorodite, stibiconite, avicennite, and various lead, zinc, and copper oxides. Unoxidized: quartz, calcite, dolomite, illite, pyrite, realgar, orpiment, stibnite, cinnabar, base- metal sulfides, and rare T1-As-Sb-S minerals.

Oxidized: quartz, clays, iron cxides, metallic gold, and calicite. Unoxidized: quartz, illite, dolomite, calcite, pyrite.

Oxidized: quartz, clays, minor calcite, metallic gold. Unoxidized: quartz, pyrite, stibnite.

Oxidized: metallic gold(?) Unoxidized: pyrite, pyrrhotite, arsenopyrite, marcasite, stibnite, orpiment, realgar, ilsemannite, cinnabar, magnetite, and metallic silver.

Oxidized: quartz, kaolinite, iron oxides, sericite, jarosite, gypsum, hexahydrite, dolomite, and calcite. Unoxidized: no published data available.

Oxidized: iron oxides, clay, quartz, metallic gold. Unoxidized: quartz, clay, cinnabar, pyrite, and arsenopyrite.

Oxidized: metallic Quartz veinlets cutting gold; 85% micron-size jasperoid. Alunite-quartz 15% coarse, visible. veinlets cutting jasperoid.

Pyrobitumen veins.

Oxidized: metallic gold. Unoxidized: gold with mercury, arsenic, tin, and thallium form thin films on pyrite. Gold is locally associated with carbonaceous matter.

Oxidized: clusters of particles between silt grains, metallic gold grains in quartz veinlets, grains in limonite pseudomorphs after pyrite. Unoxidized: associated with As and pyrite.

Quartz-calcite-orpiment, barite-galena, quartz- pyrobitumen, quartz- pyrite, and calcite veinlets. Calcite zoned away from deposit in general with quartz veins occurring within main ore deposit area.

Quartz-pyrite veinlets and post-mineralization calcite veinlets. Calcite veinlets zoned away from ore zone may have formed contemporaneous with decalcification within the ore zone.

Oxidized: metallic Quartz-stibnite and gold. Unoxidized: quartz-pyrite veins. unknown. May be with Marked increase in gold sulfides in quartz. content with silica vein-

ing and silicification.

Oxidized: no published Calcite veins in limestone data. Unoxidized: and quartz veins in phyl- metallic gold litic shale. Stockwork encapsulated by quartz veins cutting quartz. Gold also igneous rocks are in turn associated with pyrite, cut by calcite-dolomite arsenopyrite, carbon- veins. aceous matter, and magnetite. Silver associated with sulfides.

Oxidized: gold Unreported. associated with iron oxide-clay fracture coatings.

Gold is recovered from Quartz and quartz- silicified rock and magnesite veins in carbonaceous rock. silicified zones. Actual site unknown.

Klessig, 1984. Ilchik, 1984. Ainsworth and

Brimhall, 1983.

Hausen, 1967. Radtke, 1981. Radtke and

Scheiner, 1970. Hausen, 1981.

Wells, Stoiser, and Elliott, 1969.

Wells and Mullens, 1973.

Wallace and Bergwall, 1984.

William C. Bagby, U.S. Geological Survey, unpub. field notes.

Joraleman, 1951. Berger, 1975.

Wrucke, 1984. Wrucke and

Armbrustmacher, 1975.

Coppinger and Cartwright, 1983.

William C. Bagby, U.S. Geological Survey, unpub. field notes.

Table 8.2--Characteristics and form of ore in selected sediment-hosted, disseminated precious-metal deposits in the Western United states--(continued)

Alteration Ore bodies

Deposit Hypogene Supergene Form

Jerritt Canyon Decalcification, Oxidization and leaching Conformable to bedding silicification, as pods near faults remobilization of organic matter

Mercur Decalcification and Oxidization of sulfides Strongly conforms to silicification; minor to limonite; no effect bedding according to kaolinite and sericite on silicification or on variation in fracture addition clays. density, chemistry, and

proximity of faults.

Northumberland Early calc-silicate Oxidation and clay Tabular zones along skarn formation, formation tonalite sill-sediment silicification, and contact, diffuse in intense argillization breccia zones, and

stratiform bodies in sediments.

Preble

Rain

Early calc-silicate Deep oxidation along I, I t A ore zone is massive skarnification, fault zones jasperoid localized in silicification, and fault. "B" ore zone in minor argillization breccia zone of fold

hinge.

Silicification, argillization, dolomitization

Oxidation to about 200 ft

Silicification, Oxidization argillization, baritization, leaching, and oxidation(?)

Sterling -- -- Taylor

Tolman

Ore body localized along shear zone

Localized within a high angle fault zone and penetrates into wall rocks

Silicification, late Significant oxidation Tabular silicified argillization resulting in silver bodies on top of

enrichment blanket Guilmette Formation; localized by faults

Silicification followed Oxidation by late calcification

Concentrated below low angle normal fault

def ine t h e Getchel l t rend along t h e ea s t e rn flank of t h e Osgood Mountains. Rober ts (1966) included t h e Getchel l deposit in t h e Ba t t l e Mountain-Eureka trend. However, t h e location of Getchel l i s controlled by north-south s t ruc tu re s along t h e ea s t e rn flanks of t h e Osgood Mountains and i t is equivocal whether o r no t t h e Ba t t l e Mountain-Eureka bel t i s a f a c t o r in t h e location of e i ther Getchell , Pinson, o r Preble.

An impor tant a spec t of t h e regional geologic s e t t i ng of these deposits i s the i r spat ia l relationship t o

Precambrian crys ta l l ine basement and acc re t ed terranes. Ei ther t h e 0.7080 o r t h e 0.7060 s t ront ium isotope isopleth is generally accep ted a s representing t h e western e x t e n t of P recambr i an cont inenta l c rus t in Nevada (Kistler, 1983; F a r m e r and DePaolo, 1983). We have chosen t h e 0.7080 isopleth t o represent t h e western e x t e n t of P recambr i an crystall ine basement (Fig 8.2). The dashed l ine ( refer t o figure) representing t h e present-day eas ternmost e x t e n t of allochthonous western assemblage siliceous rocks


Ore bodies

Mineralogy Gold or

silver site Veins Reference

Unoxidized: pyrite, realgar, orpiment, arsenopyrite, cinnabar, and organic matter. Stibnite, barite, and quartz occur near ore bodies.

Unoxidized: pyrite, realgar, orpiment, cinnabar, barite, fluorite. Organic carbon as both kerogen and hydrocarbon fractions. Oxidized: limonite, hematite, scorodite, gypsum, melanterite, rosenite.

Oxidized: micron metallic gold, clay, iron oxides. Unoxidized: carbonaceous matter, silica, pyrite, freibergite, sphalerite, chalcopyrite, and minor molybdenite.

Oxidized: silica, goethite, lepidocrocite, hematite, sparse remnant pyrite and marcasite, and metallic gold. Kaolinite and sericite present but minor. Unoxidized: no reduced ore reported.

Oxidized: iron oxides, minor pyrite, and metallic gold. Unoxidized: pyrite, carbonaceous matter, and minor arsenopyrite, marcasite, chalcopyrite, and sphalerite.

Oxidized: barite, silica, jarosite, alunite, iron oxides, kaolinite, illite, metallic gold(?). Unoxidized: none reported.

Oxidized: iron oxides, with remnant stibnite, sphalerite, tetrahedrite, chalcopyrite, galena, pyragyrite. Unoxidized: none reported.

Unoxidized: pyrite, tetrahedrite, cinnabar, barite, calcite, quartz, gold. Oxidized: limonite, gold, same gangue mineralogy as unoxidized ore.

Gold 1-4 micron; Quartz, stibnite, and Hawkins, 1973, 1982. mineral association barite veins occur in inknown. jasperoid near ore bodies.

Calcite and arsenic minerals occur as veins in unoxidized ore.

Unoxidized: with Quartz crystals in vugs in Tafuri, 1976. pyritefmarcasite, in jasperoid, calcite-realgar Kornze and others, 1984. hydrocarbons, in veins, and orpiment-pyrite- Jewell, 1985. kerogen, and metallic marcasite-organic matter gold. Oxidized: with veins. Barite-halloysite melanterite and veins cut jasperoid. rosenite.

Quartz-silver veins of Quartz veinlets, cutting Ott, 1983. early mineralization. jasperoid. Calcite veins William C. Bagby, U.S.

cut acid-leached rock and Geological Survey, barite veins spatially unpub. field notes. associated with jasperoid.

Micron metallic gold Quartz and calcite veins Powers, 1978. occurs with As-rich cut jasperoid. Kretschmer, 1984. pyrite.

Gold encapsulated in Quartz, dolomite, Kretschmer, 1984. quartz and associated jasperoid breccia, and William C. Bagby, U.S. with pyrite. Metallic, calcite veins. Geological Survey, micron-size gold in unpub. field notes. oxidized ore.

Not reported. Quartz-alunite, quartz, Knutsen and West, 1984. barite, and calcite veins.

-- No data available.

Metallic silver. Quartz and late stage Lovering and Heyl, 1974. calcite veinlets. Havenstrite, 1983.

Gold with organic Calcite with tetrahedrite. Brady, 1984. matter and pyrite. Barite with cinnabar. Not associated with Silica micro-veinlets. tetrahedrite..

shows t h e approximate amoun t of over lap of cont inenta l c rus t by allochthonous ter ranes , beginning with t h e middle Paleozoic Antler Orogeny. Allochthonous, early-Paleozoic, western-assemblage rocks were emplaced along t h e Robe r t s Mountains thrus t in t h e middle Paleozoic during t h e Antler Orogeny (Roberts e t al., 1958). This t e r r a n e was t h e ear l ies t al lochthonous t e r r ane a c c r e t e d t o t h e western margin of t h e United Sta tes . I t was followed in turn by t h e Golconda and Sonomia t e r r anes emplaced during

t h e l a t e Paleozoic and ea r ly Mesozoic, respectively (Coney e t al., 1980). Most of t h e deposits considered in th is chapter occur within acc re t ed terranes. The except ions t o th is a r e t h e Alligator Ridge and Taylor deposits which occur e a s t of any acc re t ed terranes. Deposits within t h e Car l in and C o r t e z t rends l ie within allochthonous rocks t h a t in turn overlie Precambrian crystall ine basement. Getchel l , Pinson, and Preble in t h e Getchel l t rend l ie within acc re t ed t e r r a n e west of t h e 0.7080 strontium-isotope isopleth.

F i g u r e 8.1. L o c a t i o n s o f major sediment-hosted, d i s s e m i n a t e d p rec ious -me ta l d e p o s i t s i n the wes t - e r n U n i t e d S t a t e s . D e p o s i t s 1 th rough 3 1 are keyed to those i n T a b l e 8.5. The other deposits are: ( 3 2 ) F l o r i d a Canyon, ( 3 3 ) Weepah, ( 3 4 ) H a m i l t o n , ( 3 5 ) DeLamar, ( 3 6 ) S h a l e P i t , ( 3 7 ) Tal lman, (38) Ermont, and (39) Kendall .

The Getchel l Trend

The Getchel l trend is localized along the eas t e rn margin of t h e Osgood Mountains (Fig. 8.3). A t Getchell , in the northern par t of t he trend, mineralization is controlled by splays of t h e Getchell faul t (Berger and Taylor, 1980). I t is equivocal whether or not this faul t extends the complete distance of t h e trend. Midway along the trend, a t Pinson, t h e f au l t i s smaller and consists of a single major strand. Continuation of this s t ruc tu re south to Preble is conjecture since t h e complete length is neither exposed nor has the re been detailed geologic mapping. The re is a t Preble, however, a major north- striking f au l t t h a t controls mineralization, and t h a t may be a southern extension of the Getchell f au l t system.

Several format ions occur along the t rend and serve a s host rocks for gold, tungsten, and bar i te mineralization. These formations include the

Cambrian Preble Formation, t h e Ordovician Comus Formation, and t h e Osgood Mountains granodiorite. The Preble Formation consists of a sequence of phyllitic shales and thin-bedded t o massive l imestones (Hotz and Willden, 1964). The lower p a r t of t h e unit is in depositional con tac t with t h e lower Cambrian(?) Osgood Mountains Quartzite. This lower pa r t of t h e Preble Formation is predominantly phyllitic shale t h a t is variably calcareous and siliceous. About midway up in t h e Preble Formation the predominant lithology is limestone. The l imestone includes thin-bedded limestone separa ted by shaly partings and thick, massive turbidit ic limestone. The upper pa r t of t h e format ion is again predominantly phyllitic shale. The Preble Formation serves a s a host rock for gold mineralization a t t h e Getchell, Pinson, and Preble gold deposits and for tungsten skarn deposits around t h e margin of t he pluton.

The Ordovician Comus Formation crops ou t f rom t h e middle par t of t h e t rend north pas t t he Getchell

Figure 8.2. Locations of major sediment-hosted, disseminated precious-metal deposits in Nevada. Deposits are numbered as in Figure 8.1. The dash-dot line represents the easternmost extent of allochthonous lower Paleozoic siliceous assemblage rocks. The dashed line is the 0.7080 strontium isotope isopleth from Farmer and DePaolo (1983) and is interpreted as representing the western limit of Precambrian crystalline basement.

mine. I t consists of interbedded dolomite, l imestone, and shale with subordinate che r t , si l tstone, and tuffaceous sedimentary rocks (Hotz and Willden, 1964). Greenstones occur in t h e section near t h e Getchel l mine. This unit i s reportedly in faul t con tac t with t h e Preble Formation wherever the con tac t i s exposed (Hotz and Willden, 1964). The unit is a l t e red and mineralized a t t h e Pinson and Getchel l deposits. The rhythmic na tu re of interbedded shales and carbonate rocks is similar t o particular par ts of t h e Preble Format ion and when altered, they a r e difficult t o d i f ferent ia te particularly in the vicinity of t h e Getchel l mine.

The Cretaceous Osgood Mountain s tock in t rudes t h e Paleozoic sedimentary rocks. Con tac t - metasomat ized rocks a re host t o tungsten deposits. The pluton is predominantly granodiorite but includes

aplite, alaskite, and other der ivat ive rocks (Hotz and Willden, 1964). Cranodiorite and dac i t e dikes associated with the pluton a r e a l t e red and mineralized a t Cetchell , Pinson, and, possibly, Preble.

The s t ructura l history of t he Osgood Mountains a rea is complex. The rocks have exper ienced both the Antler and Sonoma orogenies. In addition, Erickson and Marsh (1974) identified t w o episodes of pre- Mesozoic deformation associa ted with neither the Antler nor t h e Sonoma orogeny: a L a t e Cambrian or Early Ordovician (pre-Antler) event and a L a t e Pennsylvanian or Early Permian (pre-Sonoma) event. The influence of these deformat ion even t s on the localization of t h e l a t e r sediment-hosted gold o re is ye t unclear, although t h e isoclinal folds a t t i bu ted by Erickson and Marsh (1974) t o t h e pre-Antler even t a r e important o r e controls a t Cetchell .

The southern, western, and northern (Dry Hills) margins of t h e Osgood Mountains conta in Ter t iary volcanic rocks which include basal ts and basalt ic

0 5 10


Alluvium (Quaternary and T e r l ~ a r y )

M l c m i c rocks (Quaternary and Tertiary)

Plutonic racks iCretaceous)

S e d m e n l o r y rocks (Permian l o miss is sip pion^

Cornus Formal~on (Ordovician) and Preble Formotion (Cambrian), u n d ~ v ~ d e d - - H o d lhthologies for gold

&good Mounta~n Q u o r t z ~ l e (Cambrian?)

Figure 8.3. Regional geology of the Getchell trend. Getchell, Pinson, and Preble are the major sediment-hosted, disseminated precious-metal deposits that define the trend. Geology is modified from Stewart and Carlson (1984) .

andesi tes and represent t h e youngest igneous act iv i ty in the Osgood Mountains. Quaternary basalt ic volcanism also occurred in t h e Edna Mountains a few miles southeas t of t h e Getchel l trend.

The Carlin Trend

As mentioned previously, t he Carlin t rend is identified by an alignment of gold deposits from Rain in t h e southeas t t o Dee in t h e northwest (Figs. 8.2 and 8.4). The southeas tern p a r t of t h e trend does not contain a s many identified deposits a s the northwestern part.

The Carlin t rend s t r ikes northwest and c u t s across t h e northerly t rending regional geologic fabr ic defined by eas t -di rec ted allochthonous t e r r anes (Fig. 8.2). The bel t occur s in the Rober ts Mountains allochthon which in turn, at leas t over most of the belt , overlies Precambrian crystall ine basement a s inferred from the 0.7080 strontium isotope isopleth. Of importance along th is t rend a r e uplifted windows of autochthonous rocks. These windows were the s i tes of ear ly discoveries of sediment-hosted, disseminated- gold mineralization resulting in the exploration

criterion t h a t such windows and t h e proximity of t h e Rober ts Mountains thrus t were both necessary fo r format ion of these deposits. However, many deposits have been discovered away from these windows indicating t h a t thei r presence is not necessary (e.g., Mercur, Alligator Ridge, Rain).

Rocks of t h e Rober ts Mountains allochthon a r e f rom t h e western, or siliceous, assemblage of lower Paleozoic rocks in northern Nevada (Rober ts e t al., 1958). This assemblage along t h e Carlin trend is composed predominantly of interbedded cherts, shales, and siltstones. Ca rbona te rocks a r e present but minor. These rocks a r e in th rus t con tac t with eas tern , or carbonate, assemblage lower Paleozoic rocks. The autochthonous rocks a r e predominantly silty lime- s tones and dolomites with minor shales and siltstones. Rocks in both assemblages serve a s hosts for mineralization along t h e Carlin trend. The major host format ions a r e si l ty dolomites of t he Devonian and Silurian Rober ts Mountains Formation, massive fossiliferous l imestone of a n unnamed Devonian lime- s tone unit (Evans, 19801, and che r t s and shales of t h e Ordovician Vinini Formation. These formations se rve a s hosts t o o r e in t h e northern pa r t of t he trend from

Figure 8.4. Regional geology of the Carlin trend. The major known sediment-hosted, disseminated precious-metal deposits in the trend are Rain, Maggie Creek, Gold Quarry, Carlin, Bullion Monarch, Blue Star, Gold Strike, Bootstrap, and Dee. Geology is modified from Stewart and Carlson (1984).

T r o n r l f l o n a l and aosfern oslsmbloqe rocks ( l o w s r Paieozolcl-- la A",..,.... .d pa.. - ........... roc., 0. A..l.. o........ t

Maggie Creek-Gold Quarry t o Dee t o t h e northwest. In t h e southern par t of t h e trend, t he

Mississippian Webb Format ion serves a s a host forma- t ion t o precious me ta l s a t t he Rain deposit and nearby occurrences. This format ion post da t e s the Antler Orogeny and was fo rmed from debris shed off of t he leading edge of the Ant ler highlands in to a foreland basin (Roberts et al., 1958). Shales and sil tstones a r e the predominant lithology in this formation. The upper pa r t of t h e Devonian Devils Ga te Limestone may also se rve a s a host format ion to mineralization in the southern pa r t of t h e trend. This formation may be corre la t ive with t h e unnamed Devonian l imestone t h a t serves a s a host format ion in t h e northern p a r t of t h e trend (e.g., Boots t rap and Dee).

The Carlin t r end is not only identified by an alignment of windows in the allochthon and by o re deposits, but also by Cre taceous and Ter t iary plutons. In addition, t h e t rend has positive aeromagnet ic anomalies, which, together with exposed plutons, suggest t h e presence of extensive, buried intrusions, particularly in the Car l in district . These d a t a suggest t h a t t he re is a fundamental, underlying s t ructura l control fo r t h e exis tence of this trend. As mentioned above, t h e t rend s t r ikes northwest and crosscuts the northerly trending f ab r i c defined by Paleozoic tectonism. This suggests t h a t t he Carlin t rend is younger than accre t ion of t h e Rober ts Mountains

Figure 8.5. Regional geology of the Cortez trend. Gold Acres, Cor- tez, Horse Canyon, and Tonkin Springs are the major known sediment-hosted, disseminated precious-metal deposits that occur along the trend. Geology is modified from Stewart and Carlson (1984).

allochthon. Raul Madrid (U.S. Geological Survey, oral communication, 1985) no te s t h a t folds within t h e trend a r e also northwesterly directed. He suggests t h a t t h e s t ructures t h a t define t h e t rend were developed during the Mesozoic when subduction was act ive off t h e western cont inenta l margin and plutons were emplaced e a s t of t h a t subducting margin. (Notably, the northwest-trending Miocene Oregon-Nevada l ineament (Stewart e t al., 1975) s t r ikes more northerly than the Carlin t rend and crosscuts it.) Mineralized Cretaceous igneous rocks in t h e Carlin t o Dee par t of t h e t rend suggest t h a t mineralization postdates pluton emplacement and t h a t hydrothermal fluids used the s a m e regional s t ructures t h a t served a s controls for Mesozoic pluton and dike emplacement .

Other mineralization in t h e Carlin t rend includes copper mineralization associa ted with some of t h e Mesozoic intrusions. Tungsten skarn mineralization is reported f rom Gold Str ike (Morrow and Bettles, 1982). Bar i te occurs in che r t s of t he Vinini Format ion nor th of t he Dee deposit.

The C o r t e z Trend

This t rend is defined by the alignment of t he Gold Acres, Cor tez , Horse Canyon, and Tonkin Springs deposits in north-central Nevada (Figs. 8.2 and 8.5). The t rend was identified by Rober ts (1966) a s p a r t of

GOLD A C R E S

40'15'

EXPLANATION

Ba$Olt .andesl f~ ,mnd r h y o l l f a , u n d i u l d ~ d

(ouoternorp ond ~ s r f ~ o r y l - - ~ l o r s and tuf fr

KILOMETERS

wesjern as.emalmge s l i l c s w s rocks ( Iorsr ~ a I e a z a i c ~ - -

Allochthonsus rocks o f Aof l$r oroganic bel l

T r o n r l f i o n o i and a0lf.m o s r a m b l o l e r o c k s l l o r s r Po la~ra io l - .

Autoohthonous and par . -autochthonous r o c i s at Ant le r aropanlc b a l l

t h e Ba t t l e Mountain-Eureka belt , but is now commonly referred t o a s the C o r t e z trend a f t e r t h e larges t known sediment-hosted, disseminated-gold deposit within t h e belt. More recently, Rytuba et al. (1984) have suggested t h a t Gold Acres, Cor tez , Horse Canyon, and Tonkin Springs a r e possibly r e l a t ed t o ca ldera development associa ted with the format ion of t h e Cae tano Tuff.

As with the Carlin trend, t he C o r t e z t rend contains a n alignment of mineral deposits, windows in t h e Rober ts Mountains allochthon, and Mesozoic and Ter t iary intrusive rocks. Regional host format ions include laminated sil tstones in the Devonian and Silurian Rober ts Mountains Formation, t h e Devonian Wenban Limestone, che r t s and shales of t he Ordovician Vinini Formation, and Mesozoic and Ter t iary igneous rocks. These lithologies a r e l i t t l e d i f ferent f rom those in t h e Carlin trend. Gold Acres, Cor tez , and Horse Canyon a l l occur within windows of t h e Rober ts Mountains allochthon. Tonkin Springs occurs within upper p la te rocks, although a n unmineralized Devonian l imestone, possibly lower plate, i s present. Al tered and mineralized igneous rocks within t h e t rend include Mesozoic and Ter t iary intrusions. J a m e s Rytuba (U.S. Geological Survey, ora l communication, 1985) suggests t h a t a l t e red rhyolite dikes a t Gold Acres, Cor t ez , and Horse Canyon a r e genetically re la ted t o t h e Oligocene Cae tano Tuff based on the i r t race-e lement chemistry. Silberman and McKee (1971) repor t a K/Ar a g e of 94.3 i. 1.9 m.y. on ser ic i te f rom a n a l t e red and mineralized fe ls i te sill in t h e Gold Acres mine. This age implies a Cre taceous age for t h e mineralization. The resolution of t h e ac tua l a g e of t h e Gold Acres deposit must awa i t fur ther geochronological studies.

Other mineralization in the C o r t e z t rend includes the C o r t e z silver deposit in t h e Cambrian Hamburg Dolomite, t h e Buckhorn gold deposit in Pliocene volcanic rocks, and bedded ba r i t e in Ordovician sedimentary rocks. The C o r t e z silver deposit i s a replacement deposit of mantos and fissure veins in t h e Hamburg Dolomite (Gilluly and Masursky, 1965; Gilluly and Gates, 1965). The Buckhorn deposit is an epi thermal hot-spring re la ted vein deposit.

Isolated Deposits

Several of t he sediment-hosted, disseminated precious-metal deposits in Nevada occur a s deposits isolated from other known deposits of th is type. However, many of these isolated deposits do occur in recognized mineral bel ts t h a t a r e defined by regional s t ruc tu res and t h e alignment of several d i f ferent types of mineral deposits. Examples of isolated deposits a r e Taylor, Northumberland, and Alligator Ridge.

The Taylor deposit occurs in t h e eas tern pa r t of t h e s t a t e (Fig. 8.2) and is located in t h e east-west or iented Hamilton-Ely bel t of Rober t s (1966). This bel t is defined by a n alignment of Ter t iary intrusions, a positive aeromagnet ic anomaly, and o re deposits (Stewar t et al., 1977). The o r e deposits in th is bel t include several d i f ferent types; however, t he Taylor mine is t h e only known major sediment-hosted, disseminated precious-metal deposit. There a r e severa l prospects near Taylor t h a t may turn o u t t o be deposits of this type.

The regional geology of t h e Taylor deposit is character ized by Paleozoic sedimentary rocks intruded by Ter t iary rhyolit ic rocks. The sedimentary rocks a r e predominantly l imestone, dolomite, and shale. The Devonian Gui lmet te Format ion serves a s t h e host rock fo r t h e silver deposits a t Taylor. This is a widespread formation and also hosts silver o re at Hamilton, 80 km west of Taylor. The Gui lmet te Format ion is transit ional upward with t h e Devonian t o Mississippian Pilot Shale. Although t h e Pilot Shale is unmineralized in t h e Taylor area , i t serves a s t h e major host for gold o r e a t Alligator Ridge (see below). The overlying Mississippian Joana Limestone hosts minor silver o r e in t h e Taylor d is t r ic t and hosts major sulfide replacement o re a t Ward, 20 km west of Taylor. The Joana Limestone is overlain by t h e Mississippian Chainman Shale and the Mississippian t o Permian Ely Limestone. These units s e rve a s host t o copper o r e a t t h e Ruth porphyry copper deposit and t h e Ely Limestone also hosts lead, silver, and z inc o r e a t Ward.

Northumberland (Fig. 8.2) i s l oca ted along t h e eas t e rn margin of t h e Northumberland caldera. The deposit i s located 40 km nor theas t of t he volcanic- hosted Round Mountain gold deposit (Tingley and Berger, 1985) and lies 60 km nor theas t of t h e Manhattan silver-gold dis t r ic t (Roberts, 1966; Shawe and Stewar t , 1976). Although other sediment-hosted deposits occur south of Northumberland (e.g., t h e Shale P i t near Round Mountain and t h e White Caps mine in the Manhattan district) , t h e location of t h e Northumberland deposit s eems t o be closely re la ted t o ca ldera development and a s such, i s a truly isolated deposit of this type.

The regional geology of t h e Northumberland a r e a is character ized by lower Paleozoic rocks intruded by Mesozoic granodiorite followed by format ion of the middle Ter t iary Northumberland caldera (McKee, 1974). The Paleozoic rocks consist of both autochthonous and allochthonous rocks of t h e Antler Orogeny (Kleinhampl and Ziony, 1984). The allochthonous rocks consist of cher t , shale, argillite, and some interbedded greenstones, including pillow basalts. These a r e considered t h e Ordovician Vinini Format ion by McKee (1974). Autochthonous rocks consist of l imestone, shale, and argi l l i te mapped a s t h e Ordovician Pogonip Group and t h e Silurian Masket Shale. Gold and silver o res occur in the Mesozoic intrusions and bedded bar i te occurs in t h e allochthonous rocks. Silver-rich qua r t z veins occur in Cre taceous tonal i te bordering t h e Northumberland deposit. The silver mineralization is apparently older than t h e sediment-hosted gold.

Alligator Ridge (Fig. 8.2) i s located northwest of Ely, Nevada, and appears unrela ted t o any known mineral belts. I t occurs about 20 km south of t he Bald Mountain mining district . This d is t r ic t contains l imestone replacement o re bodies, qua r t z vein and contact -metamorphic deposits, and placer deposits (Hose e t al., 1976).

The regional s t ra t igraphy in t h e Alligator Ridge a r e a includes a Devonian through Mississippian section overlain by Ter t iary volcanic rocks. The Devonian pa r t of t h e section consists of t h e Devils G a t e Limestone and t h e lower pa r t of t h e Devonian t o Mississippian Pilot Shale. These format ions a r e

W. C. BAGBY & B. R. BERGER 183

overlain by t h e Mississippian Joana Limestone, Chainman Shale, and Diamond Peak Formation. This is the s a m e s t ra t igraphic section t h a t occurs in t h e Taylor district (see above). Alligator Ridge is one of the few deposits of this type t h a t ac tual ly has some geologic character is t ics t h a t a r e suggestive of shallow emplacement of t h e o r e (e.g., possible hydrothermal breccias).

GEOLOGIC CHARACTERISTICS O F THREE END-MEMBER, SEDIMENT-HOSTED, DISSEMINATED

PRECIOUS-METAL DEPOSITS

The Carlin, Preble, and Taylor deposits a r e representa t ive of t he end-member ca tegor ies t h a t we recognize in this deposit type. Carlin serves t o i l lus t ra te the ca tegory t h a t is character ized by fine- grained, disseminated-gold o r e t h a t is difficult t o visually sepa ra t e f rom unaltered rock. Preble i l lustrates t h e ca tegory t h a t is character ized by gold ore associated with in tense sil icification (jasperoids) and quar tz veining. Both of these deposits a r e gold- rich end-members. Taylor i l lustrates t h e silver-rich end-member and is also character ized by in tense silicification in t h e o re zone. There is obviously a gradation between t h e chemical and geological character is t ics of these end-members. We feel, however, t h a t they serve a s easily recognized examples of t he diversity in this deposit type. The geologic character is t ics of these th ree deposits a r e reviewed here t o provide a basis for comparison of t h e deposits l isted in Tables 8.1 and 8.2 and a summary diagram of t h e character is t ics t h a t define the end- members i s given in Table 8.3.

Carlin

The Carlin deposit (Figs. 8.2, 8.4, and 8.6) is probably t h e most extensively studied deposit of this type. The recognition of this deposit a s possibly a new type of gold deposit (Hausen, 1967) influenced exploration models and resulted in government,

academic, and industry research in to t h e genesis of t h e deposit. The following deposit character is t ics a r e summarized from Hausen (19671, Hausen and Kerr (19681, Radtke et al. (1980)) Hausen (19811, Hausen e t al. (19831, Radtke (19811, and Adkins and R o t a (1984).

Host lithology--Most o r e a t Carlin occurs in the upper pa r t of t h e Devonian and Silurian Rober ts Mountains Formation. The most favorable host rocks in th i s sect ion a r e laminated, argillaceous, arenaceous dolomites and calcareous mudstones. Arenaceous peloid wackestone a lso occurs in this section, but is apparently less favorable a s a host for ore.

Igneous rocks--Dikes along f au l t s a r e exposed in t h e Carlin pit. Although these dikes a r e hydrothermally a l tered, thei r presumed original composition was felsic, including possibly granodiorite and qua r t z diorite. Small s tocks a r e exposed nor th of Carlin t h a t vary in composition f rom granodiorite t o diorite. The stock a t t h e Cold Str ike mine has a K-Ar d a t e from bioti te of 121 m.y. (Hausen, 1967). Bioti te from a dike in t h e southwestern p a r t of t h e Carlin P i t yielded a d a t e of 131 m.y. (Morton et al., 1977). This l a t t e r d a t e should be used with caution a s several unpublished da te s f rom a l t e red rocks in t h e Carlin mine yield K-Ar ages t h a t vary f rom l a t e Ter t iary t o Mesozoic (M. L. Silberman, U.S.G.S., personal communication, 1985).

Structure--The Rober ts Mountains thrus t faul t is exposed in t h e Carlin p i t (Fig. 8.6). This faul t brings western siliceous assemblage rocks over eas tern carbonate f ac i e s rocks, both of ear ly Paleozoic age. The thrus t faul t i s in turn disrupted by s e t s of normal faults. These f au l t s s t r ike N. 40-45 E., N. 25-30 E., and N. 25 W., in order from oldest t o youngest. Both upper and lower p l a t e rocks a r e folded with fold axes striking N. 45 W.

Alteration and oxidation--There has been both hypogene and supergene a l tera t ion of t h e rocks a t t h e Carlin deposit. Supergene a l tera t ion (oxidation and acid-leaching) was so in tense in some a reas t h a t i t is difficult t o in t e rp re t ear l ier hypogene alteration. Alteration associa ted with primary mineralization can only be understood by examining unoxidized, ore-

Table 8.3--Comparison table summarizing geologic characteristics of jasperoidal and pod-like, sediment-hosted, disseminated precious-metal deposits

Jasperoidal, quartz- veinlet type

Disseminated, pod-like type

Quartz veins common Quartz veins are uncommon

Main ore type is in silicified rock Main ore type is not silicified

Ore primarily restricted to fault zones Pod-like ore bodies extend away from faults

Several different silicification stages Jasperoid may be present

Both gold- and silver-rich varieties Gold-rich variety most common

Siliceous rocks common Calcareous rocks common


F e i s ~ c d i k e s ( ~ r e t a c e o u s ~ ~

Al lochthonou~ rocks of Antler orogenlc belt

0 100 2 0 0 u

METERS

Autochthonous rocks of Antler orogen lc b e l t -

Figure 8.6. Plan map of the early (1972) pit at the Carlin deposit. The stippled pattern represents combined oxidized and unoxidized ore zones within the deposit. Geology is modified from Adkins and Rota (1984).

Dsrm Roberts Mountolnr Format ion U (D ....... and S,i".,."I

Comblned axldlzed and unoxldized are ~n t h e 1 9 7 2 pit

P i t outline

,----- C o n t a c t

)4L. R o b e r t s Mountains thrust

,--- H l g h onple f a u l t s

3 Strike and dip of bedding

bearing, and barren rocks. Hypogene a l tera t ion consists of severa l types: (1) decalcification; (2) argillization; (3) silicification; and (4) calcification. Decalcification was ear ly and resulted in the removal of carbonate minerals from t h e silty dolomites and limestones. Ca lc i t e was preferentially leached over dolomite resulting in increased porosity and permeability. Argillization accompanied decalcification resulting in format ion of i l l i te f rom de t r i t a l feldspar with minor amounts of montmoril lonite and kaolinite. Sil icification was a bulk replacement process and resulted in jasperoid formation. L a t e calcification resulted in t h e format ion of ca l c i t e veins.

Hypogene a l tera t ion zoning of t h e overall deposit is clearly defined by calc i te and sil ica distribution. Ca lc i t e i s sparse in t h e main ore zone but is abundant above and away f rom the ore horizon. Silica occurs beneath t h e o r e horizon and is localized along faults.

Supergene a l tera t ion includes oxidation of sulfide minerals and acid-leaching of t h e rocks. This resulted in the format ion of iron, arsenic, and antimony oxides and sul fa tes and in t h e redistribution of calcite. Supergene oxidation is deepest along major faul ts and format ion contacts.

Ore types--Different ore types a t Carlin have been described by Hausen (1967) and Rad tke (1981); t he following review is based on thei r information. The o re zones a t Carlin a r e localized near faul ts (Fig. 8.6). The o r e is commonly difficult , if not impossible, t o visually sepa ra t e from waste. Ore control is s t r ic t ly dependent upon gold assay, and severa l types of o r e a r e recognized based on metallurgical behavior. Rad tke (1981) recognized five primary unoxidized o re types: (1) normal; (2) siliceous; ( 3 ) carbonaceous; (4) pyritic; and ( 5 ) arsenical.

Normal o r e is composed of dolomite, illite, and quartz. I t contains minor amounts of kaolinite, ser ic i te , and some remnant calcite. Py r i t e i s t h e most abundant sulfide and occurs a s remnant d iagenet ic subhedral cubes and two types of introduced pyrite: dispersed cubes sca t t e red in the rock . m a t r i x and c lus ters of small (<IOU) framboidal pyrite. The origin of t h e framboidal pyr i te i s controversial and is in terpre ted by Hausen (1981) a s having formed from bacteria. Gold, in normal ore, coats sulfides and a small amoun t occurs a s meta l l ic gold associa ted with quartz.

Siliceous o r e is less than 10% of the deposit and

W. C . BAGBY & B . R. BERGER 185

i s cha rac te r i zed by sil ica replacement of carbonate. This o re conta ins 80-95% quartz, 5-10% clay (illite, ser ic i te , and kaolinite), 1-5% dolomite, and small amounts of organic carbon. Gold in this type of o re occurs a s me ta l l i c gold par t ic les ( < l o p ) encapsulated in hydrothermal quartz.

Carbonaceous o res a r e those t h a t contain enough organic carbon t o inhibit t h e complexing of gold with cyanide. Organic carbon in th is o re type varies from 1-6% and, excep t f o r th is high carbon content, carbonaceous o r e resembles normal ore. Gold is apparently associated with sulfides, and Radtke (1981) did not observe any meta l l ic gold in th is o r e type.

Pyr i t ic o r e contains 3-10% pyr i te compared t o t h e 0.5-3% pyr i te con ten t of o ther unoxidized o re types. Py r i t i c o r e commonly occurs a s bands cut t ing carbonaceous ore. The pyr i te occurs a s subhedral t o euhedral grains up t o 0.4 mm in diameter with lesser amounts of framboidal pyrite. The matr ix is composed of de t r i t a l qua r t z grains, fine-grained dolomite, and hydrothermal quartz. Detr i ta l dolomite is corroded and l i t t l e remnant ca l c i t e is present. Most of t h e gold in th is o re type occurs a s coatings on framboidal and cubic pyrite.

Arsenical o re i s character ized by 0.5-10% arsenic, mostly in t h e fo rm of realgar with some orpiment and a f ew sulfosalts. The o re i s composed of silt-sized qua r t z grains in a matr ix of dolomite grains, clay, and carbonaceous material . Fossil f ragments a r e silicified. Arsenic minerals occur a s crosscutting veins and fillings of small (50 -300~) vugs. Cold occurs a s fi lms on both framboidal and cubic pyrite.

In addition t o pyrite, o ther minor sulfides and sulfosalts occur a t Carlin including realgar, orpiment, st ibnite, cinnabar, te t rahedr i te , tennant i te , sphalerite, galena, molybdenite, chalcopyrite, chalcocite, covelli te, lorandite (TIAsS2), chr is t i te (TIHgAsS3), weissbergite (TISbS2), ,ellisite (TI3 Ass3), and car l in i te (TIZS). Nat ive arsenlc has also been reported a t Carlin.

Oxidized o re is composed of varying amounts of quartz, clay (illite, with lesser amounts of kaolinite, ser ic i te , and montmorillonite), and dolomite. Oxidized o re contains only about 0.03-0.354 organic carbon and is thus t a n o r bleached. Limonite staining is common. The oxidized o r e is most likely formed by supergene oxidation of the primary sulfide-bearing ores. Rad tke (1981) and Hausen (19671, however, both suggest t h a t late-stage oxidation and acid leaching resulting f rom boiling of a geothermal system a t and near the surface was responsible for oxidation and acid leaching a t Carlin. Although boiling of hydrothermal waters can occur a t depths t h a t a r e typical of t he porphyry environment (Cunningham, 19781, Radtke (19811, and Hausen (1967) appeal t o a model t h a t requires the deposit t o have formed in e i ther t h e near-surface or t h e surface (hot-spring) environment. An a l ternat ive hypothesis derived from cur ren t research by C. A. Kuehn a t Carlin ( refer t o Bodnar e t al., this volume) is t h a t t h e format ion of gold o re a t Carlin occurred in a reducing environment a t depth ( a t leas t one kilometer) and t h a t t h e r e were no hypogene oxidizing fluids (with respect t o pyrite). In light of this interpretation, a l l oxidized o r e a t Carlin formed from supergene a l tera t ion of t h e primary ores a t low temperature .

T race elements--Harris and Rad tke (1976) investigated t h e geochemistry of unoxidized o re a t Carlin. Their in tent was t o t e s t t h e correlation of gold with mercury, arsenic, and antimony t h a t had already been established at other deposits (e.g., Getchel l (Joralemon, 1951); Gold Acres (Wrucke and Armbrustmacher, 1975)). In addition, they s ta t i s t ica l ly analyzed t h e distribution of gold, barium, copper, molybdenum, lead, zinc, boron, tellurium, selenium, and tungsten in t h e west, main, and e a s t o re zones. They found high correlations between gold, arsenic, mercury, and antimony. Gold is negatively corre la ted with barium, and barium has relatively high corre la t ions with the base metals. Cold and tellurium have a significant positive correlation t h a t i s s t rongest in t h e e a s t o r e zone.

Organic geochemistry--Much of t h e o r e in t h e main and e a s t o re zones is carbonaceous. Hausen and Ker r (1968) noted t h a t concentra tes t h a t a r e rich in carbon commonly a r e enriched in gold. However, Hausen (1967) notes t h a t t h e presence of carbonaceous ma te r i a l is no a priori evidence of ore. A similar gold- carbon correlation was noted by Rad tke and Scheiner (1970). The difficulty has been in defining t h e type of gold-carbon association and whether or not t he carbonaceous mater ia l served a s a transporting or precipitating agent fo r gold. Radtke and Scheiner (1970) theorized severa l organic compounds t h a t could se rve t o che la t e with gold. Unfortunately, until t h e chemis t ry of t he hydrothermal solutions responsible for gold t ranspor t and deposition a r e clearly understood, t h e theoret ica l gold-carbon chela tes can not be suitably tested.

Vein types--The deposit i s known fo r i t s lack of veins. However, d i f ferent types of veins do in f a c t occur, but thei r spatial and genet ic association with gold i s not completely understood. Vein types include bar i te , ca lc i te , quartz, and quartz-pyrobitumen veins. Hausen (1967, p. 73) shows photomicrographs of gold associated with quar tz veins. This indicates t h a t a t l ea s t some of t h e qua r t z veining occurred during gold deposition. Unfortunately, due t o the fine-grained na tu re of t h e gold, i t has been difficult fo r subsequent workers t o identify t h e gold-quartz vein association. For example, quartz-pyrobitumen veins may have formed prior t o gold deposition even though many of them now occur in ore. Barite veins a r e in terpre ted by Hausen (1967) t o be earlier than t h e gold stage. The s ta t i s t ica l analysis of barium, base metals, and gold by Harr is and Radtke (1976) indicate a negative corre la t ion between gold and barium t h a t possibly suppor ts t he geologic interpretation of the t w o having been deposited a t d i f ferent times. Ca lc i t e veins, possibly formed during gold deposition, occur in zones peripheral t o silicified zones and l a t e r ca l c i t e crosscuts acid-leached, oxidized rocks.

Paragenesis--Development of a deta i led paragenet ic sequence in this o re deposit is ext remely difficult due t o t h e general lack of crosscutting veins t h a t can unequivocally be re la ted t o gold or o ther a spec t s of hydrothermal activity. Two paragenet ic sequences have, however, been published. Radtke (1981) in terpre ted four stages: (1) ear ly and main t h a t included the introduction of quartz, pyrite, potassium- bearing clays, hydrocarbons, minor barite, and ear ly

jasperoid; (2) l a t e main with introduction of As, Sb, Hg, TI, S, Au, and l a t e r base me ta l minerals; (3) an acid-leaching and oxidation s t age with t h e introduction of major barite, jasperoid, anhydrite, and ca l c i t e veins; and, (4) l a t e oxidation caused by weathering. On t h e o the r hand, Hausen (1967) placed ba r i t e and t h e base me ta l s very early (Cretaceous) i n t h e paragenesis and gold, realgar, st ibnite, na t ive arsenic, jordanite, tennant i te , cinnabar, and jarosite midway (Tertiary) in t h e paragenesis with la ter , probably Pliocene, supergene oxidation. The major d i f ferences between these two paragenet ic in terpre ta t ions a r e t h e p lacement of t he base-metal and bar i te veins and t h e length of hydrothermal activity. We fee l t h a t Cretaceous mineralization is predominantly base me ta l and not associated with t h e gold o r e a t Carlin. In addition, although t h e oxidized o re a t Carlin probably did form by supergene oxidation, t h e t iming of th is is probably l a t e Miocene t o early Pliocene, a known period of deep weathering throughout Nevada.

Fluid inclusions--Radtke e t al. (1980) repor ted fluid-inclusion analyses by John Slack (U.S. Geological Survey, referenced in Radtke , 1981) on quartz, barite, realgar, and sphalerite f rom t h e Carlin deposit. The inclusion d a t a were in terpre ted in t e r m s of Radtke's paragenesis and indicated main s t age (4 samples) temperatures of 200°C increasing t o about 3 0 0 ' ~ during t h e acid-leaching s t age (13 samples). Car l A. Kuehn (U.S. Geological Survey, unpublished data , 1985) i s currently studying fluid inclusions at Carlin and has noticed some important compositional character is t ics of inclusions in qua r t z (Kuehn and Bodnar, 1984).

EXPLANATION

Grovel (Quaternary)

Stable isotopes--Radtke e t al. (1980) published stable-isotope analyses fo r t h e deposit. They analyzed oxygen, hydrogen, carbon, and sulfur isotopes in veins, rocks, minerals, and fluid inclusions. Their resul ts indicate t h a t hydrothermal solutions were highly exchanged me teo r i c waters with 6D between -140 and -160 per mil. The 9 4 s values of 4.2-16.1 per mil were in terpre ted a s indicating diagenet ic pyrite a s the source of sulfur in t h e deposit.

Age of mineralizatibn- he a g e of mineralization a t Carlin has not yet been di rect ly dated. An igneous dike t h a t i s a l t e red and mineralized within t h e deposit was dated a t 131 t 4 m.y. (bioti te) by Morton et al. (1977). The Gold Str ike pluton north of Carlin is a l tered, contains gold, and has been da ted a t both 78.4 + 3.9 m.y. (biotite, Morton et al., 1977) and 121 * 5 m.y. (biotite, Hausen, 1967). Rhyolit ic lavas and domes west of t he Carlin p i t over l ie jasperoid and a r e not a l tered or mineralized. These lavas a r e da ted a t 14.2 * 0.3 m.y. (sanidine, McKee et al., 1971). Thus, t h e age of t h e deposit could be a s old a s 130 m.y. but no younger than 14 m.y.

Taylor

The Taylor deposit in eas t e rn Nevada (Figs. 8.2 and 8.7) is t he silver-rich end-member in our classification of sediment-hosted disseminated precious-metal deposits. In addition, Taylor has character is t ics t h a t a r e more similar t o t h e jasperoid- qua r t z vein type (Preble) t han t h e disseminated type (Carlin). The Taylor deposit has not been extensively

Lattte ond docite flows and tuffs (Tertlory)

Rhyolite dikes (Ter t iary or Mesozoic)

Eiy Limestone (Permian to Misstrs~pp~onl

Choinman Shole ( M ~ s s ~ s s l p p ~ a n )

0 Joana Limestone ( M l s s ~ s s ~ p p i a n )

P ~ i o t Shole (Missaslppian and Devonian)

Gu~lmet te Formation (Devonian) - - D i v ~ d e d lnto

Member b Member a

Dolomite and (Devonian)

Figure 8.7. Geology of the Taylor mining district. Jasperoids are spatially associated with faults. The large jasperoid near the center of the Figure is the present site of the Taylor deposit. Geology simplified from Drewes ( 1 9 6 2 , 1967 1.

a Slllclfied rock (jasperold)

- M E T E R S

Thrust f a u l l Teeth on upper p o t e 10

Y Strlke ond d # p of b e d $

studied and, thus, published information i s scarce. > T h e following summary abou t t h e Taylor deposit is derived f rom Havenst r i te (19831, Hose e t al. (19761, Lovering and Heyl (19741, Drewes (1962, 19671, J e f f r ey M. Edwards, (Newmont Exploration Limi ted , unpublished data , 19851, and personal observations (William C. Bagby, U.S. Geological Survey, unpublished field notes, 1985).

Host lithology--The Middle and Upper Devonian Cu i lme t t e Format ion i s t h e host of silver o r e a t Taylor. This format ion is composed of four members a s mapped by Drewes (1967). Member a is a thick- bedded l imestone t h a t i s locally replaced by a massive coarse-grained dolomite. Member b is shaly l imestone and coarse-grained dolomite t h a t in ter tongues with, and grades la tera l ly into, limestone. Member c includes l imestone and shaly l imestone with a sand- s tone marker bed nea r i t s base. This member is t h e major host of ore. These t h r e e members together comprise t h e lower p a r t of t h e formation. Member d is reef l imestone, shaly l imestone, sandstone, and conglomerate t h a t is only locally preserved.

Igneous rocks--Rhyolite dikes in t rude Paleozoic and Mesozoic rocks in t h e Taylor distr ict . These dikes s t r ike north and no r theas t and a r e hydrothermally a l t e r ed t o clay. Notably, severa l of t h e dikes contain xenoliths of jasperoid, t h e predominant o re t ype a t Taylor (Havenstri te, 1983). The dikes have been mapped a s Mesozoic o r Te r t i a ry (Drewes, 1962). Havenst r i te (1983) repor ts a d a t e of 35 m.y. fo r t h e dikes, but does not indica te how t h e d a t e was determined.

Structure--Drewes (1962, 1967) mapped normal f au l t s trending north, nor thwest , and nor theas t in t h e Taylor distr ict . These f au l t s c u t t h e rhyolite dikes and a r e thus considered a s l a t e Te r t i a ry t o Holocene f au l t s associa ted with Basin and Range tectonism. Havenst r i te (1983) i n t e rp re t s folding in t h e Taylor d is t r ic t a s a la rge over turned ant i form. The axis of t h e ant i form is remarkably coincident with o n e of Drewes's (1962) normal f au l t s t h a t Lovering and Heyl (1974) r e f e r t o a s t h e Taylor fault . Havenst r i te (1983) suggests t h a t t h e o r e bodies a t Taylor formed in crackle breccias t h a t developed in t h e hinge and along the flanks of t h e antiform. Lovering and Heyl (1974) and Drewes (1967) suggest ins tead t h a t t h e o r e bodies were formed in and near f au l t breccia zones and t h a t t h e normal f au l t s controlled hydrothermal fluid flow in t h e district .

Al tera t ion and oxidation--Silicification and argil l ization a r e t h e only t w o a l t e r a t ion types t h a t a r e reasonably well documented in t h e Taylor distr ict . Thin-bedded and shaly l imestones, par t iculary beneath thick shale horizons, a r e commonly silicified near f au l t s (Drewes, 1962, 1967). The si l icif ication is a replacement process result ing in t h e format ion of jasperoid. O r e bodies in t h e Taylor d is t r ic t a r e l imited t o t hese jasperoid zones. Argil l ization in t h e d is t r ic t includes t h e a l tera t ion of fe ldspar in t h e rhyolite dikes t o kaolinite (Drewes, 1962, 1967). The dikes a r e notably unmineralized. Although sulfides s t i l l occur a t t h e su r f ace overing and Heyl, 19741, t h e Taylor deposit is well oxidized (Havenstri te, 1983). This oxidation remobilized silver, t hus result ing in supergene enr ichment of t h e ore. Acids fo rmed during supergene

a l tera t ion of sulfides impar t ed a leached, bleached, and spongy appearence t o t h e su r f ace jasperoids.

O r e types--Silver-bearing jasperoid is t h e predominant o r e t y p e a t Taylor. Jasperoids formed f rom limestone along and away f rom fau l t s a s though t h e f au l t s served a s conduits fo r hydrothermal solutions (Lovering and Heyl, 1974). The o r e bodies occur in t h e upper p a r t (member c) of t h e Gui lmet te Format ion and a r e tabular bodies generally about 50 f e e t th ick (Havenstri te, 1983). Sulfides associa ted with t h e si l icif ication a r e s t ibni te , sphaleri te, te t rahedr i te , chalcopyrite, pyrite, galena, and pyrargyr i te (Lovering and Heyl, 1974). Gangue minerals include ca lc i te , dolomite, ser ic i te , tourmaline, monazi te , bari te, f luorite, and a p a t i t e (Lovering and Heyl, 1974). Limoni te is a common consi tuent of ore. Matrix qua r t z in highly mineralized jasperoid conta ins minute inclusions of organic carbon, st ibnite, limonite, cerargyr i te , and ant imony oxides (Lovering and Heyl, 1974). Silver occurs a s pyrargyr i te (?), cerargyr i te , a rgent i te , and meta l l ic si lver f inely disseminated throughout t h e jasperoid.

T race elements--Lovering and Heyl (1974) investigated t h e geochemis t ry of jasperoids in t h e Taylor d is t r ic t and compared them t o unmineralized jasperoids f rom elsewhere in Nevada. They noted t h a t Ag, Au, Cu, Hg, Pb, Sb, Te, and Zn a r e a l l anomalous in Taylor jasperoids compared t o unmineralized jasperoid (Lovering, 1972).

Organic geochemistry--The only mention of organic carbon in any of t h e published l i t e r a tu re on Taylor is by Lovering and Heyl (19741, who no te t h e presence of minute inclusions of carbon in jasperoid quartz. Havenst r i te (1983) appeals t o t h e organic-rich Chainman Shale a s a sou rce rock fo r both silver and silica introduced in to t h e Gu i lme t t e Formation.

Vein types--Veins of ca l c i t e , quar tz , and bar i te occur in t h e Taylor distr ict . Bar i te veins a r e zoned peripheral t o t h e Taylor deposit (J. M. Edwards, Newmont Exploration Limi ted , ora l communication, 1985) whereas both ca l c i t e a n d qua r t z veins cross c u t jasperoid a t t h e deposit. I t i s equivocal whether o r not any of t hese veins a r e pa r t of t h e s a m e mineralizing even t t h a t deposited silver.

Paragenesis--Drewes (19671, Lovering and Heyl (1974), and Havenst r i te (1983) a l l a g r e e t h a t si lver and si l ica were deposited simultaneously f rom hydro- t he rma l solutions flowing through fault-breccia zones. Havenst r i te (1983) indica tes t h a t rhyoli te dikes and sills must b e younger t han t h e o r e deposit ing episode s ince xenoliths of jasperoid occur in t h e dikes. Havenst r i te (1983) c i t e s th is a s evidence t h a t magmat i c f luids were not associa ted with t h e hydro- t he rma l sys tem a s proposed by Drewes (1962, 1967) and Lovering and Heyl (1974). However, Lovering and Heyl (1974) m a k e a s t rong a rgumen t t h a t t h e dikes and t h e hydrothermal fluids used t h e s a m e s t ruc tu re s a s pathways t o t h e upper crust . I t i s ent i re ly possible t h a t t h e dikes a r e simply l a t e intrusive s t ages of a much larger magmat ic-hydrothermal sys tem t h a t had been deposit ing si lver fo r considerable t i m e prior t o dike emplacement .

Fluid inclusions--Jeffrey M. Edwards (Newmont Exploration Limi ted , ora l communication, 1985) has performed preliminary heat ing studies on quartz,

fluorite, and la te-s tage sphaler i te from the Taylor district . He in t e rp re t s boiling fluids during deposition of t h e qua r t z and f luor i te at average temperatures of 2 0 0 ' ~ and 220°c, respectively. Inclusions in late- s t age sphalerite (filling vugs and coat ing jasperoid f r agmen t s in breccias) t h a t contain th ree phases a t room tempera tu re and homogenize a t about 300°C a r e in terpre ted a s C 0 2 - H 2 0 inclusions.

Age of mineral~zation--The o r e deposits in t h e Taylor d is t r ic t have not been di rect ly dated. If t h e rhyolite dikes were emplaced about 35 m.y. ago, a r e unmineralized, and conta in f r agmen t s of jasperoid, then they place an upper l imit on t h e a g e of mineralization. A lower a g e l imit is provided by t h e age of t h e host rocks which is middle and l a t e Devonian. The nor th and northwest-trending ore- bearing faul ts in t h e d is t r ic t c u t Ter t iary volcanic rocks nor th of t h e d is t r ic t (Lovering and Heyl, 19741, indicating a Ter t iary age fo r t h e faulting. This places a t ighter const ra in t on t h e mineralization a g e by l imiting i t t o t h e Tertiary.

Preble

The Preble deposit is t he southernmost and smal les t known deposit of th is type in the Getchell trend. I t serves t o i l lus t ra te the end-member ca tegory of gold-rich deposits t h a t contain in tense silicification and relatively common qua r t z veining in the o r e zone (Figs. 8.2 and 8.8). Mineralized rock can a lmost a lways be visually d i f ferent ia ted f rom unmineralized rock due t o a l tera t ion and qua r t z veining. The following reveiw is summarized from Kretschmer (1984) and William C. Bagby and Raul J. Madrid (U.S. Geological Survey, unpublished data , 1985).

Host lithology--The Cambrian Preble Format ion is t h e host rock for t h e deposit. The format ion consists of phyllitic shale with interbedded, finely laminated and massive limestones. The phyllitic shales host most of t he ore. These rocks, where f resh and unaltered, a r e composed of variable amounts of

muscovite, quartz, chlorite, carbonate, and smect i te . Metamorphic, folded qua r t z veins occur throughout the phyllitic shale sequence. Quar t z over growths on these veins contain gold and a r e evidence of l a t e r silica addition associated with gold deposition.

Igneous rocks--An a l t e red fe ls ic dike, striking northerly and dipping t o t h e e a s t cu t s through the deposit. The dike occupies a faul t zone t h a t had r ecu r ren t movement post-emplacement of t he dike. The genet ic relationship of t he dike t o t h e gold is unknown.

Structure--The sequence is folded and faulted and crenulation c leavage is well developed in t h e phyllitic shale. A t l eas t t w o periods of deformation have a f f ec t ed t h e rocks. Early, e i ther pre- or syn- metamorphic veins a r e folded. These veins consist of ca l c i t e where they c u t l imestone and quar tz+calc i te where they c u t phyllitic shale. North- and eas t - trending high-angle f au l t s offse t t h e sequence. The f au l t t h a t contains the fe ls ic dike is one of the main controls of t h e geometry of t h e ore body. The north- striking faul ts have dropped the Preble Formation down t o t h e east .

Alteration and oxidation--Alteration of t h e phyllitic shale involves an addition of silica and a change in color. This color change is character ized by bleaching of the original phyllite and t h e greenish sheen becomes buff. Dolomitization of the l imestones in t h e sequence is common near high-angle faults. Jasperoid veins c u t t he l imestone and sil ica replacement of l imestone forms massive jasperoid. Carbonaceous m a t t e r has been locally remobilized in f au l t zones.

O r e types--Ore a t t h e Preble deposit is associa ted with some form of silicification, e i ther a s replacement of l imestone or phyllitic shale (jasperoid) or a s quar tz veining. The quar tz veining is bes t observed in thin sect ion and under t h e scanning e l ec t ron microscope a s quar tz overgrowths on metamorphic quar tz veins and a s patchy replacement of host rock. The main ore zone a t Preble is localized

O X I D I Z E D .'\-


Preble Formotion (C0rnbrlan)--Conslsls of

Figure 8.8. Schematic cross section of the main ore zone of the Preble deposit. The ore horizon is restricted to a shear zone associated with an east- dipping normal fault. The diagram is modified from Kretsch- mer (1984).

c-.-c.--. O Y I I ~ ~ S S zone o f s o l d ore - F o Y I I

within a north-striking faul t zone. Mineralization is fa i r ly uniform within t he zone and is predominantly l imi ted t o t h e f au l t and hanging-wall host rocks. Both oxidized and unoxidized o re occur. Gold occurs with sulfides and iron oxides a f t e r pyrite. The ave rage s i ze of t h e gold gra ins i s about 2 microns (Kretschmer , 1984). Pyr i te , both cubic and framboidal var ie t ies , is t h e most abundant sulfide with arsenopyrite, marcas i te , chalcopyr i te , and sphaler i te a s minor accessory sulfides (Kretschmer, 1984).

Trace-e lement geochemistry--Crone (1982) and Crone et al. (1984) inves t iga ted t h e geochemistry of iron oxide-rich f r a c t u r e coatings over t h e deposit. Their results i nd i ca t e a s t rong geochemical corre la t ion of gold with mercury , arsenic, and antimony. Kre t schmer (1984) repor ts a correlation of thallium, barium, and f luor ine wi th gold.

Organic geochemistry--Unoxidized o r e occurs in black, carbonaceous, phyllitic shale. The carbonaceous m a t t e r has been hea t ed t o high t empera tu re s and is a lmost graphite. There i s no positive corre la t ion between abundance of organic carbon and gold (W. C. Bagby, U.S.G.S., unpublished data , 1985).

Vein types--The Preble deposit i s a stockwork veinlet , jasperoidal deposit. There a r e severa l generations of veining a t Preble, some of which a r e definitely associa ted with t he period of gold deposition. Vein types include metamorphic quar tz , quartz-carbonate, and ca l c i t e veins, and l a t e r quar tz , quar tz-ca lc i te , jasperoid, dolomite, and ca l c i t e veins. These veins a r e a l l recognizable in t h e field with t h e possible exception of d i f ferent ia t ing between metamorphic and gold-related qua r t z veins. The jasperoid veins cu t t i ng l imestones a r e presumably t h e veins sampled by Crone (1982) t o identify t h e gold, arsenic, mercury, ant imony association.

Fluid inclusions--Fluid-inclusion studies a r e now in progress on Preble samples (W. C. Bagby, unpublished da t a , 1985). Metamorphic quar tz- ca rbona te veins conta in numerous t ra i l s of secondary inclusions. These include both two phase (liquid and vapor) and t h r e e phase (liquid C 0 2 , liquid H20 , and vapor) inclusions a t room t empera tu re t h a t a r e generally less t han 10 microns across.

Isotopes--No stable-isotope d a t a a r e available f o r t h e deposit.

Age of mineralization--M. L. Silberman (U.S. Geological Survey) col lec ted a l t e r ed fe ls ic dike ma te r i a l f rom t h e mine l ized zone. Ser ic i te f rom t h e ld dike yielded a n 4 0 ~ r / Ar a g e of 100.42 k1.6 m.y. (L.W. Snee, U.S. Geological Survey, wr i t t en communication, 1980). The in terpre ta t ion is t h a t Preble is a similar a g e t o Getchel l (Berger, 1980).

GENERAL ASPECTS O F TRACE ELEMENT AND STABLE-ISOTOPE GEOCHEMISTRY

In a n a t t e m p t t o compare t race-e lement concentra t ions between deposits, we chose 0.30 ppm Au a s a cut-off for se lec t ing samples f rom t h e published l i te ra ture . In addition, we chose t o compare only Ba, Hg, Sb, As, Ag, and Au s ince these e l emen t s a r e generally repor ted in t h e l i t e r a tu re fo r deposits of th is type. In th is way, we hoped t o be able t o compare

t h e t race-e lement con ten t of rocks ranging f rom low- grade o r e (0.30 pprn) t o higher grades. The published l i t e r a tu re commonly has only s ta t i s t ica l summaries of d a t a wi thout tabula t ions of a l l data. Where th is was t h e c a s e (Carlin, Getchell , and Gold Acres) we used t h e ranges reported, even though in a l l t h r e e cases;the ranges included gold values less t han 0.30 ppm gold. Trace-e lement geochemis t ry is given in Table 8.4 and summarized in Figure 8.9.

The highest gold values repor ted a r e those f rom Car l in and Cor tez . The Car l in samples represent both oxidized and unoxidized o r e f rom within t h e early, shallow pit (Radtke et al., 1972) whereas t h e C o r t e z samples a r e a l l su r f ace samples col lec ted prior t o mining (Erickson e t al., 1966). Although Rad tke et al. (1972) show t h a t oxidized o r e has a range of gold

Figure 8.9. Summary of trace element variations in sediment-hosted, disseminated precious- metal deposits. The trace element data are restricted to samples containing 0.30 ppm Au unless otherwise noted in the text or in Table 8.4. The dash-dot extensions to the ranges for Pinson and Preble indicate higher anomalies in iron-oxide fracture coatings than those in rocks (Crone, 1982; Crone et al., 1984). The dashed line extensions with arrows indicate that some values were either greater than or less than the detection limits of the instru- ment.

Table 8.4--Summary of trace-element geochemistry for selected sediment-hosted, disseminated precious-metal deposits

Deposit Au Ag As Sb Hg Ba (Reference) Rock type (# ) (#) (#) (il) ( # ) (#)

Bootstrap Altered dike 0.30-2.0 0.5-10.0 200-500 100-500 1.1-3.0 200-300 (Evans, 1974) (2) (2) <(2) <(2) (2) (2)

Siltstone 0.4 2.0 <ZOO 200 0.26 150 (carbonate assemblage) (1) (1) (1) (1) (1) (1)

Gouge, siliceous breccia 1-14 0.7-3.0 <ZOO-1,000 100-1,000 0.14-8.0 700->5,000 (siliceous assemblage) (3) (3) (3) (3) (3) (3)

Siliceous breccia 0.3 0.5-20 <200 <loo-5,000 0.8-1.5 300-700 (carbonate assemblage) (2) (2) (2) (2) (2) (2)

Carlin Limestone Ore (Radtke and (unoxidized) others, 1972)

Oxidized Ore

Cortez Limes tone 0.34-116 tr.-5.5 <500-5,000 <loo-150 .02-2.7 NA (Erickson, (carbonate assemblage) (11) (11) (11) (11) (11) et al., 1966)

Getchell Carbonaceous ore 0.99-3.91 NA 2,800-48,000 NA 45.8-641.8 NA (Berger, 1975) (4) (4) (4)

Siliceous ore 0.51-4.49 NA 57,000-128,000 NA 248.9-400.1 NA (2) (2) ( 2 )

Argillic ore 3.09-6.00 NA 64,000-78,000 NA 89.3-93.9 N A (2) (2) (2)

Gold Acres Limestone and dolomite <0.02-30 <0.5-100 <200->10,000 <loo-500 <0.1->10 20-5,000 (Wrucke and (137) (137) (137) (137) (137) (137) Armbrus tmacher, 1975)

values from 0.04-125 ppm compared with 0.02-88 ppm in unoxidized ore, t he medians a r e 11 and 10 ppm, respectively. With respect t o d i f ferences in t race- e lement concentrations between unoxidized and oxidized o r e a t Carlin, Ag, As, and Ba medians a r e t h e same, whereas the Hg median is lower, and the Sb median is higher in t h e oxidized rocks (Radtke et al., 1972). Silver values a r e generally less than about 10 ppm in gold-rich members of t h e deposit type, but do reach a s high a s 100 ppm (Gold Acres). The silver-rich

end member , Taylor, has high gold values of about 10 ppm and high silver values of about 2,000 ppm. In general, for t hese deposits, As varies from 100 t o 1,000 ppm, Sb f rom 5 t o about 200 ppm, Hg f rom 0.2 t o about 30 ppm, and Ba f rom 30 t o about 1,000 ppm. Although the re a r e exceptions t o these ranges, t he re is considerable overlap between t h e di f ferent deposits. The amount of overlap in t h e concentra t ions of these e lements indicates t h a t t h e chemical and physical character is t ics of t h e hydrothermal sys tems were


Notes

LLD: Au (0.05), Ag (0.5), As (200), Sb (loo), Hg (0.05), Ba (20.0), Cu (5.0), Pb (10.0), Zn (200.0). Eleven samples analyzed; 3 had 0.05-0.15 ppm Au, two had less than detection limit, and 4 had no Au detected.

LLD: Au (0.05), Ag (0.5), As (ZOO), Sb (loo), Hg (0.05), Ba (20.0), Cu (5.01, Pb (10.0), Zn (200.0). Eleven samples; four had 0.05-0.1 ppm Au; six had no Au detected.

LLD: AU (0.05), Ag (0.5), As (zoo), Sb (loo), ~g (0.05), Ba (20.0), Cu (5.0) Pb (10-0), Zn (200.0). Three samples total.

LLD: Au (0.05), Ag (0.5), As (200), Sb (loo), ~g (0.05), Ba (20.0), Cu (5.0), Pb (10.0), Zn (200.0). Two samples analyzed. All "jasperoids9' analyzed had Au at 0.1 ppm.

LLD: Au (0.02), Ag (0.7), AS (lo), Sb (0.5), Hg (0.01), Ba (lo), T1 (50), Cu (I), Pb (7), Zn (5). Sample suite included 120 unoxidized and 250 oxidized samples. LLD: Au (0.02), Ag (0.7), As (lo), Sb (0.5), Hg (0.01), Ba (lo), T1 (50), Cu ( I ) , Pb (7), Zn (5). Most from upper beds of Roberts Mountains Formation. Samples from pit exposures and drill cuttings.

LLD: Au ( ? ) , Ag (?I, As (500), Sb (loo), Hg ( ? ) . Surface samples from Cortez district. Detection limits not recorded.

LLD: Au ( ? ) , As ( ? ) , Hg (? ) . Six analyses of carbonaceous ore reported. Two had 0.21 ppm. No detection limits reported.

LLD: Au ( ? ) , As ( ? ) , Hg (?) . Three analyses of siliceous ore reported. One had 0.21 ppm Au. No detection limits reported.

LLD: Au ( ? ) , As ( ? ) , Hg (? ) . Two analyses of argillic ore reported. No detection limits reported.

LLD: Au (0.02), Ag (0.5), As (Zoo), Sb (loo), Hg (.01), Ba (201, Cu (5), Pb (lo), Zn (200). Samples collected within Gold Acres pit.

similar, even between gold- and silver-rich systems. However, depar ture f rom t h e norm f o r severa l of t h e deposits indica tes t h a t t h e hydrothermal sys t ems were each unique; possibly ref lec t ing in tens i ty of t h e sys tem ( tempera ture , hydrologic flow ra t e , l i fe t ime) and also possibly t h e chemis t ry of t h e source and host rocks. A more comprehensive in terpre ta t ion is current ly no t possible due t o t h e incompleteness of t h e comparat ive d a t a base.

Stable-isotope s tudies of th is deposit t ype a r e

even more incomplete t han t race-e lement investigations. R y e et al. (1974) and Rad tke e t al. (1980) inves t iga ted stable-isotope sys temat ics a t C o r t e z and Carlin, respectively. R y e (1985) reviewed t h e conclusions of t h e preceding papers and added addit ional d a t a f rom o the r deposi t s fo r comparison. Figure 8.10 shows t h e ranges in 6 3 4 ~ f o r sulfides and sul fa tes f rom di f ferent deposits. F o r t h e most par t , t h e bar i tes have sulfur isotopic s ignatures similar t o sedimentary barite. On t h e o the r hand, sulfides f rom these deposits

Table 8.4--Summary of trace-element geochemistry for selected sediment-hosted, disseminated precious-metal deposits--(continued)

Deposit Au Ag As Sb Hg Ba (Reference) Rock type ({I) (#) (#) ({I) ({I) ( # )

Pinson Jasperoid 0.34-34.29 NA 245-1,650 0.2-74.8 N A N A (Powers, 1978) ("A" ore zone) (60) (60) (60)

(Crone, 1982) Rock (lithology 0.30-1.65 <0.1-0.3 190-1,000 6.0-41.0 2.10-16.00 NA unrecorded) (6) (6) (6) (6) (6)

Iron oxide fracture 0.34-5.24 <0.05-1.51 69-9,000 2.0-317 0.38-42 N A coatings (8) (8) (8) (8) (8)

Preble Rock (lithology 0.95-1.7 0.1-0.4 240-530 13-34 6.6-16 NA unrecorded) (4) (4) (4) (4) (6)

Iron oxide fracture 0.31-4.73 <0.05-0.80 53-13,140 21.1-330 0.44-24 N A coatings (10) (10) (10) (10) (10)

(William C. Shale (drill chips) 0.31-5.7 0.08-15 21-1,200 4.3-210 0.15-75 NA Bagby, U.S. (28) (28) (28) (28) (28) Geological Survey, unpublished data.

Taylor Jasperoid (Lovering and Heyl, 1974)

Windfall Hamburg Dolomite (Grove, 1979)

Dunderburg Shale

Andesite

NA, not analyzed; ND, not detected; and #, number of samples analyzed; all analyses are in ppm; LLD, lower limit of detection.

have isotopic s ignatures t h a t overlap with sedimentary although C o r t e z and Carlin may be of d i f ferent age , and organic sulfur. Detailed in terpre ta t ions of t hese t h e hydrothermal sys t ems were very similar. variations a r e p rema tu re given t h e lack of de ta i led paragenet ic sequences fo r t h e samples.

Oxygen, hydrogen, and carbon isotopes have been SUMMARY O F GEOLOGIC CHARACTERISTICS examined a t Car l in and C o r t e z (Rye, in press). La rge negat ive 6D values indica te dominance of t h e fluids by Sediment-hosted, disseminated precious-metal me teo r i c water. Large 6180 values indica te deposits display considerable var ia t ion in the i r subs tant ia l exchange between fluids and rock in both geologic character is t ics (Tables 8.1 and 8.2). We have deposits. R y e (1985) uses t hese d a t a t o sugges t t h a t t r i ed t o examine th is variation by defining two end-

Notes

LLD: Au ( ? ) , Ag ( ? ) , As ( ? ) , Sb (? ) . Sixty-one samples analyzed. One had 144 pprn Au (not included in tabulation). Detection limits not listed.

LLD: Au (0 .05) , Ag (0 .1 ) , As (1 .0 ) , Sb ( 5 . 0 ) , Hg (0 .01) . Surface samples, lithology unrecorded. Thirty samples analyzed, 22 had no detectable Au; 2 had 0.05 pprn Au.

LLD: Au (0 .01) , Ag (0 .05) , As (1.01, Sb (2 .0 ) , Hg ( 0 . 0 5 ) , Cu ( 2 0 ) , Zn ( 5 ) . Thirty samples, same site as rocks; 22 samples had 0.01-0.30 pprn Au. Twenty-two soil samples; 4 had 0.05-0.10 pprn Au, all others less than 0.05 pprn Au.

LLD: Au (0 .05) , Ag ( 0 . 1 ) , As (1 .0 ) , Sb ( 5 . 0 ) , Hg (0 .01) . Thirty-two samples collected; 5 had 0.05- 0.25 pprn Au, 23 had <0.05 pprn Au.

LLD: Au (0 .01) , Ag (0 .05) , As (1 .0 ) , Sb ( 2 . 0 ) , Hg (0 .05) , Ba (0 .00) , T1 (0 .00) , CU ( 2 0 ) , Pb (0 .00) , Zn ( 5 ) . Thirty-two samples; 22 had 0.01-0.26 pprn Au.

LLD: Au (0 .05) , Ag (0 .01) , As (0 .1 ) , Sb ( 0 . 0 2 ) , Hg ( 0 . 0 1 ) , T1 (0.01) . One hundred five samples; 54 had <0.05 pprn Au, 22 had 0.05-0.29 ppm Au, one had 31 pprn Au. All 20 ft composites.

LLD: Au ( ? ) , Ag ( ? ) , As ( ? ) , Sb ( ? ) , Hg ( ? ) , Ba ( ? ) , Cu (?I, Pb ( ? ) , Zn (? ) . Three "representative samplest' from Taylor district. No detection limits reported.

LLD: Au ( 0 . 1 ) , As ( l o o ) , Sb ( 1 , 0 0 0 ) , Pb (101 , Zn (10) . LLD: Au (0.11, As ( l o o ) , Sb ( 1 , 0 0 0 ) , Pb ( l o ) , Zn (10) . High Au value of 23 ppm not included. A total of 92 samples were analyzed for these elements. Thirty-four samples had 0.1-0.2 pprn Au, 13 had trace, and 18 had nil. Fourteen shale samples analyzed; 5 not analyzed for Au, 3 had 0.01-0.2 pprn Au, 2 had tr. Au, and 1 had nil.

LLD: Au ( 0 . 1 ) , As ( l o o ) , Sb ( 1 , 0 0 0 ) , Pb ( l o ) , Zn (10) . Three andesite samples analyzed; one had 0.01 pprn Au, other two had >0.3 pprn Au.

member subsets t h a t represent e x t r e m e cases of t h e geologic variation of this deposit type. But in addition, t he re a r e a large number of similarit ies t h a t provide cohesiveness and definition t o t h e deposit type. Together, t h e variations and similarit ies provide a basic geologic framework from which i t i s possible t o compare known deposits and t o develop exploration and assessment programs. One shortcoming of th is approach is t h a t i t assumes a similar level of understanding for the deposits used in defining t h e

deposit type. This is simply not t h e case. However, t h e current in teres t in this deposit t ype is result ing in increased industry, academic , and governmental research t h a t will expand our knowledge of these deposits.

Regional and District Scale

The regional and dis t r ic t geologic character is t ics seem t o show t h e leas t variation. Whether t h e deposit

I I I I I I

I S U L F I D E S

I BARITE

Figure 810. Summary of sulfur isotopic variation in sulfides and barite from several sediment-hosted, disseminated precious- metal deposits. The range given for sulfides includes data for pyrite, stibnite, realgar, and orpiment from Carlin, Cortez, Mercur, Jerritt Canyon, White Caps, and Getchell. Barite range includes data for samples from Carlin, Cortez, Getchell, Jerritt Canyon, Mercur, Pinson, Preble, and White Caps. Other ranges are: (A) igneous sulfur: (B) sulfide sulfur from the Roberts Mountains Formation; (C) organic sulfur: and (D) sedimentary sulfate. Data are from Rye (1985), Guenther (1973), and Ohmoto and Rye (1979).

is of t he disseminated, pod-like o r jasperoidal, quartz- veinlet subset and whether i t is gold- or silver-rich is not dependent a t our present level of knowledge upon a cer ta in regional geologic environment. Sedimentary rocks of any a g e t h a t include thinly bedded silty or shaly carbonate rocks or ca lcareous sil tstones or shales provide t h e ideal host environment. Fluid pathways formed by regional faul t ing and folding se rve t o concen t ra t e and di rect hydrothermal gold-bearing solutions. A regional h e a t source is also needed t o drive a mineral-depositing hydrothermal system. I t may be t h a t regional lithology chemically controls whether o r not a deposit i s gold- o r silver-rich. The d a t a base t o support th is i s incomplete.

Deposit Scale

There a r e common threads t h a t run through the di f ferent deposits.

Host lithology--Thinly bedded sil ty or shaly, carbonaceous l imestones and dolomites, calcareous, carbonaceous shales and sil tstones, and bioclastic l imestones a r e t h e most common types of host lithologies.

Igneous rocks--Intermediate t o si l icic dikes, plugs, domes, and (or) s tocks a r e present a t most deposits; e i ther in the deposit itself or within the district . These rocks range f rom Cre taceous t o middle Tertiary. The genet ic association between these rocks and the gold mineralization is unknown.

Structure--High-angle normal and (or) strike-slip f au l t s a r e present in a l l deposits. These f au l t s a r e generally pre-, syn- and post-ore. Thrust faul ts a r e present a t some deposits, but do no t seem t o b e t h e major controll ing s t ruc tu re for o r e deposition. Folds a r e present, both regionally and within deposits. In f ac t , many deposits occur near t h e c r e s t of regional antiforms. The deposit-scale folding in most deposits has not been mapped in detail and, thus, t he control of t hese s t ruc tu res on o r e deposition is unknown.

Al tera t ion and oxidation--Hypogene a l tera t ion a lmost invariably includes ear ly decalcification followed by silicification. The disseminated, pod-like deposits contain silicified o re in and near faults, but much of t h e o re in th is subtype is not strongly silicified enough t o be t e rmed a jasperoid. The jasperoidal,, quartz-veinlet subtype deposits contain mostly s i l ~ c ~ f i e d ore. Other a l tera t ion includes format ion of clays and local remobilization of carbonaceous ma t t e r . All deposits a r e oxidized t o some extent . Oxidized o r e is usually bleached, contains iron oxides and sulfates, and may be acid leached.

Ore types, shapes, and mineralogy--ore in the jasperoidal, quartz-veinlet subtype is generally res t r ic ted t o a f au l t zone with only minor leakage in to t h e surrounding wallrock. On t h e other hand, disseminated, pod-like deposits characterist ically contain o re zones t h a t a r e podiform and extend away f rom fau l t zones. The mineralogy of e i ther subtype always includes pyrite. Other common, but variably present, minerals include cinnabar, st ibnite, arsenopyrite, f luorite, barite, ca lc i te , and various thallium and arsenic sulfides and sulfosalts. Oxidized and unoxidized ores a r e present in both subtypes. However, in some of the smaller jasperoidal, quar tz- veinlet deposits, only t h e oxidized ore is mined.

Si te of gold--The gold s i t e has not been identified in a l l deposits. However, where i t is known, t h e gold par t ic les a r e always submicroscopic, on t h e order of 1 t o 5 microns. Metallurgical t e s t s indicate t h a t gold is associatied with pyrite, clay, silica, and organic mat ter . The na tu re of these associations is not well understood.

Trace-element geochemistry--This deposit type is well known fo r t h e t race-e lement association of Au, As, Sb, Hg, and TI. These e lements a r e a lmost universally present in deposits of this type, although concentra t ions a r e variable. Other e lements t h a t a r e anomalous in some deposits include W, Te, Se, Cd, and F.

Organic geochemistry--Carbonaceous m a t t e r i s present in some form in t h e unoxidized o r e of a l l known deposits. This includes amorphous carbonaceous ma t t e r , pyrobitumen, graphite, and kerogen. Hydrocarbons were identified a t Mercur (Tafuri, 19761, but a r e a s ye t unknown in o ther deposits. The genet ic relationship between gold deposition and carbonaceous m a t t e r , if any, i s unknown.

Vein types--Veins a r e most abundant and the i r relationship t o gold seem the c leares t in t h e jasperoidal, quartz-veinlet subset of deposits. However, veins a r e also present, and possibly very important , in the disseminated, pod-like deposits. Veining typically includes qua r t z and calc i te with variable barite, f luorite, and dolomite veins.

W. C. BAGBY & B. R. BERGER 195

Para~enes i s - -T he general lack of c lear crosscut t ing relations of d i f ferent veins and thei r temporal association with gold inhibits t he development of detailed paragenet ic sequences in these deposits. However, t he general paragenesis a lmost always involves early decalcification and carbon remobilization with l a t e r silicification and, in some cases, very l a t e ca l c i t e veining. Supergene oxidation and acid-leaching a r e common, but of variable intensity, a t a l l deposits.

Isotopes--Only t h e Cor tez , Carlin, and Mercur deposits have been examined in deta i l in t e rms of stable-isotope systematics. These da ta suggest t he importance of me teo r i c fluids in t h e formation of the deposits and sedimentary sulfur a s t h e source of sulfide sulfur. More reconnaissance and detailed stable-isotope studies a r e necessary t o understand the variation in this deposit type.

Age of mineralization--There is evidence t h a t these deposits formed a s ear ly a s Cretaceous and a s l a t e a s middle Tertiary. Getchel l i s t he only deposit t h a t has been da ted by radiometr ic techniques resulting in a Cre taceous age for mineralization. Al tered and mineralized Ter t iary dikes a t o ther deposits indicate a Ter t iary age for mineraliztion. I t is unlikely t h a t t he re is a sediment-hosted, disseminateed precious-metal epoch of mineralization. Instead, i t is more likely t h a t t he re i s a spread in mineralization ages.

ENVIRONMENT OF FORMATION

Deposits of th is type have been classified a s epi thermal t o t e l e the rma l (near-surface) deposits since t h e f i r s t descriptions of Getchell (Joralemon, 1951) and Carlin (Hausen, 1967). Joralemon (1951) c i t ed the telescoped na tu re (meaning t h e close association of gold, realgar, st ibnite, and cinnabar over a short ver t ica l distance) of t h e mineralization, large vugs and an intensely sha t t e red o re zone a s evidence fo r epi thermal deposition a t Getchell. Hausen (1967) discussed t h e difficulty of classifying Carlin due t o t h e lack of vuggy, banded veins t h a t a r e character is t ic of t he epi thermal environment. He noted, however, t h a t silicified zones a t Carlin had the appearance of siliceous sinter and thus chose t o classify Carlin a s "low-temperature epi thermal bordering on telethermal." Both authors referred t o Graton's (1933) coinage of te le thermal a s t h e extension of epi thermal into the near-surface t o hot-spring environment. Jora lemon (1951) drew correlations between cinnabar- bearing, hot-spring deposits and mercury a t Getchel l whereas Hausen (1967) drew analogies between t race- e lement associations (Hg, Sb, Au, and Ag) a t Carlin and Steamboat Ho t Springs, Nevada.

This deposit type has thus been considered epi thermal to possibly te le thermal without any rigorous pressure and t empera tu re studies. The epi thermal environment was described by Lindgren (1933) a s a subset of his genet ic classification of hydrothermal o r e deposits. He suggested t h a t epi thermal deposits a r e formed by ascending hot waters (hydrothermal) of uncertain origin in a shallow environment t h a t i s dominated by rapid changes in

both h e a t and pressure. Lindgren placed t empera tu re constraints of between 5 0 ~ - 2 0 0 ~ C at moderate pressures t o define t h e epithermal-hydrothermal environment of ore deposition. A modified Lindgren classification in terpre ts "moderate pressures" a s between 40-240 bars a t depths of 150-915 me te r s (Ridge, 1968). The t e l e the rma l environment a s defined by Graton (1933) is t h e terminous of hydrothermal ac t iv i ty and thus, can be in terpre ted t o represent hot- spring activity. Although Graton (1933) did not assign any depth or t empera tu re constraints t o t h e t e l e the rma l environment, presumably i t would be the cooler and shallower ends of Lindgren's (1933) epi thermal environment.

Research in progress on sediment-hosted, disseminated precious-metal deposits suggests t h a t t h e environment of deposition for such deposits may in f a c t ac tual ly be deeper than previously assumed, extending depths of format ion t o deep epi thermal and mesothermal zones. The mesothermal environment a s defined by Lindgren (1933) is typified by t empera tu res of 2 0 0 ~ - 3 0 0 ~ C a t high pressures ( in terpre ted by Ridge (1968) t o represent depths of one to th ree kilometers). Various geological cr i ter ia observed in these deposits suggest t h a t t h e deeper environment may be t h e case for these deposits. In f ac t , Lindgren (1933, p. 563) classified Mercur, U tah , and the sediment-hosted, disseminated-gold deposits in t h e Moccasin dis t r ic t in Montana a s mesothermal. I t i s c lear t h a t a complete understanding of t h e genesis of deposits of this type is s t i l l elusive and waits t h e completion of cu r ren t research.

EXPLORATION APPLICATION

The di f ferences and similarit ies in geologic character is i tcs between deposits of this t ype provide a wealth of recognition c r i t e r i a for deposits of this type. The regional character is i t ics (host lithology, s t ructure , and heat) , summarized above, a r e t h r e e c r i t e r i a t h a t may be used t o identify possible a reas favorable fo r exploration. Of course, t h e alignment of deposits of this type in recognized mineral bel ts i s probably t h e f i rs t and most important cr i ter ion used on the regional scale by explorationists.

District- and deposit-scale cr i ter ia include s t ructure , a l tera t ion, and geochemical character is t ics . These c r i t e r i a a r e necessarily based upon deta i led geologic mapping. For example, a l tera t ion types and thei r spat la l relationships t o s t ruc tu res and di f ferent potential host lithologies mus t be mapped and reasonably understood prior t o geochemical sampling. In addition, vein types and thei r crosscutting relationships help def ine a t a r g e t area. Most of t h e deposits discussed above have l a t e ca l c i t e veining crosscutting oxidized rocks. Jasperoidal breccia and jasperoid veins generally occur near ore , even when they themselves may not ca r ry high gold values. Geochemical surveys, including rock and soil, can be ext remely valuable for closing in on a t a r g e t (Bagby et al., 1984). The ubiquitous gold sui te of associa ted t r a c e e lements , arsenic, mercury, and ant imony is a n important indicator of gold mineralized rock. Extremely high values for these e l emen t s a r e

no t necessary t o define a favorable area. Instead, i t may be more significant t h a t t h e sui te of indicator e l emen t s i s present.

INFLUENCE OF GEOLOGIC CHARACTERISTICS ON MINING

Grade and Tonnage

Sediment-hosted, disseminated precious-metal deposits a r e variable in grade and tonnage. Table 8.5 l ists 31 gold- and silver-rich deposits, 30 for which the re a r e grade and tonnage data. Although the deposits a r e located throughout the western United Sta tes , most a r e from Nevada. Grade and tonnage values for these deposits a r e shown on cumulative- frequency graphs in Figure 8.11. Gold grade is plotted against tonnage for 29 of the deposits in Figure 8.12. The cumulative-frequency diagrams show t h a t t h e median gold grade is 2.5 gramst tonne and the median tonnage is 5.1 me t r i c tonnes. The gold curve also

CARBONATE-HOSTED GOLD-SILVER

10 1 I , I # . , , I , nz35

TONNES ( I N MILLIONS)

a.

indicates t h a t 90% of t h e deposits have gold grades less than 7.6 gramsltonne. The larges t deposits in t e rms of gold grade within this sui te a r e Carlin and J e r r i t t Canyon. The cumulative-frequency tonnage curve indicates t h a t 90% of t h e deposits have less than 24 million tonnes of ore. The only deposit with greater tonnage is Gold Quarry. A cornparison of individual deposits is best displayed by the grade versus tonnage plot in Figure 8.12. The diagonal lines across the figure indicate to ta l contained grams (or ounces) of gold for any given grade and tonnage. This shows t h a t t he larges t deposits in t e r m s of contained gold a r e Carlin, Getchell , J e r r i t t Canyon, and Gold Quarry. All of t he o ther deposits contain less than 50 million grams (about 1.5 million t roy ounces) of contained gold per deposit.

One in terpre ta t ion t h a t can be derived from the grade and tonnage d a t a i s t h a t t h e chances of finding deposits greater than 24 million tonnes a re small. Of t h e sample suite, t h e r e a r e really only 4 outstanding deposits, in t e rms of contained gold, out of 29. Of those four, two (50%) were discovered during the

CARBONATE-HOSTED GOLD-SILVER 10 , , , , , , I , , . \ n i 3 5

a

0 0 0016 0063 016 3 4 10 2 5 6 3 6 4 0 100

GOLD GRADE ( IN g / l )

b

0 4 10 2 5 6 3 16 40 100 250 630 1600 4000

SILVER GRADE ( IN g / i )

C

CARBONATE-HOSTED GOLD-SILVER

Figure €3.11. Cumulative-frequency diagrams for grade and tonnage of sediment-hosted, disseminated precious-metal deposits. Some data used in these diagrams are con£ idential and thus, the total number of deposits (n = 35) is greater than those listed in Table €3.5. Figures are from Bagby and Singer (1983).

10

0 9 -

2 o e - - Y) 0 0 7 -

W n 0 6 -

LL 0 0 5 -

Z p 0 4 - + g 0 3 - n : 0 2 - O O 1 -

0 0

, , , , , , , , , n 35

\-• - - - - - - - - - 15

1 ( 4 I . . , 8 8

Million Tonnes

F i g u r e 8.12. Gold g r a d e v e r s u s tonnage f o r gold-r ich , sediment-hosted, d i s s e m i n a t e d p rec ious -me ta l d e p o s i t s . T h e d i a g o n a l l i n e s i n d i c a t e to ta l c o n t a i n e d gold f o r a g i v e n grade a n d t o n n a g e . D e p o s i t a b b r e v i a t i o n s are those u s e d i n T a b l e 8.5. (Go ld p r o d u c t i o n f o r 1 9 8 4 i n the U n i t e d S t a t e s w a s 2.3 m i l l i o n troy o u n c e s (71.5 m i l l i o n g r a m s ) as e s t i m a t e d b y the U.S. B u r e a u o f Mines, M i n e r a l Commodity Summaries, 1985).

recent increase of exploration act iv i ty in t h e 1970's. Each of t h e remaining more recently discovered and developed deposits contain less than 50 million grams of gold. The explorationist , however, should not be discouraged. The f a c t t h a t 50% of these outstanding large deposits were found in recent years i s reason enough t o hope t h a t more will be discovered, especially in light of our increasing knowledge of t h e geologic character is t ics of this deposit type.

Mineability

Deposits of this t ype a r e minable due t o t h e coincidence of several geologic and economic factors. The continued high marke t value for precious meta ls , particulary in a n otherwise difficult me ta l s economic environment, is t he pr ime reason the exploration for, and development of, low-grade precious-metal deposits continues. The high-market value of precious meta ls (particularly gold) balanced against mining costs and me ta l ext ract ion is favorable. Geologic f ac to r s t h a t influence minability a r e oxidation, large tonnages near t h e surface , and the fine-grained, disseminated na tu re of t h e gold. Most of

these deposits a r e deeply oxidized. The oxidation has essentially performed t h e f i r s t metallurgical process by f ree ing the gold f rom i t s association with e i ther sulfide or carbonaceous mater ia l . This oxidized o r e is therefore amenable t o heap-leaching, making deposits with grades a s low a s 1 gramttonne profitable t o mine. Early re turns on oxidized ore can also provide capi ta l for development of c i rcui ts necessary t o recover meta ls f rom the unoxidized o re t h a t is encountered in t h e l a t e r s t ages of mining. Since these deposits a r e near surface and the gold is character is t ica l ly evenly distributed in i t s host, t h e deposits a r e bulk minable. This lowers mining costs which a r e important f ac to r s in deriving a profit f rom this type of deposit.

The economic balance sheet for these deposits is important t o industry and t h e favorable figures result in increased capi ta l outlays for exploration and development. This deposit t ype is therefore presently of in teres t t o t h e mining industry in t h e United States. Geologic studies of individual deposits and comparisons between and within deposits increase our knowledge of thei r occurrence and genesis, and this knowledge leads t o fur ther discoveries.

Table 8.5--Grade and tonnage data for sediment-hosted, disseminated precious-metal deposits

Tonnage Au grade Ag grade (million)

Location troy ozIgram troy oz/gram short tonlmetric ton Notes

1. Alligator Ridge White Pine Co., 0.09913.39 Unknown 5.915.4 1983 reserves plus production (AL) NV to date, company handout.

2. Atlanta Lincoln Co., 0.0812.74 1.6154.86 1.111.0 Bonham (1983). (AT) NV

3. Bald Mountain Lawrence Co., 0.19916.82 0.372112.75 7.616.9 Norton (1983) Bald Mountain ( BA) SD district; several mines.

4. Blue Star Eureka Co., 0.1214.11 Unknown 1.811.6 Bonham (1983). (BL) NV

5. Bootstrap Eureka and Elko 0.04911.68 Present 1.411.3 McQuiston and Shoemaker (1981) (BO) Co., NV Calculated from 717,000 tons

shipped to Carlin at assumed grade .064 oz; 165,000 tons leach at .044 oz; 500,000 tons leach at .028 oz.

6. Candelaria Mineral Co., 0.00610.21 41137.14 514.5 Bonham (1983) Bagby (1983, (CN) NV unpublished field notes).

7. Carlin Eureka Co., 0.32/10.97 Unknown 11/10 Jackson (1983). (CAI NV

8. Cortez Lander Co. , 0.27919.57 Unknown 3.613.3 Bonham (1983) Actually (CO) NV 3,562,100 tons.

9. Dee Elko Co., 0.0913.09 Present 413.63 Wallace (in press). (DE) NV

10. Getchell Humboldt Co., 0.19416.65 Unknown 15.4113.97 Bonham (1983) Total (GE) NV production, reserves, and

possible.

11. Gold Acres Lander Co., 0.06512.23 Unknown 2.812.54 Bonham (1983). (GO) NV

12. Gold Quarry Eureka Co., 0.04311.47 Unknown 1831166 Skillings (1984). (GQ) NV

13. Gold Strike Eureka Co., Unknown Unknown Unknown Bonham (1983) Greater than (GS) NV 100,000 oz Au.

14. Hilltop Lander Co., 0.079/2.71 Unknown 5.114.6 Bonham (1983). (HI) NV

15. Horse Canyon Eureka Co., 0.1013.43 Unknown 5.014.5 Bagby (1983, unpublished field (HO) NV notes).

16. Jerritt Canyon Elko Co., 0.24318.33 Unknown 14.06112.76 Hawkins (1982), Anonymous (JE) NV (1982).

17. Maggie Creek Eureka Co., 0.092/3.15 Unknown 4.814.35 Anonymous (1980). (MA) NV

18. Manhattan Nye Co., 0.03611.23 Unknown 5.014.54 Bonham (1983). (MN) NV

19. Mercur Tooele Co., 0.10213.56 Unknown 14.31112.98 Anonymous (1981), Larimer (ME) UT (1983).

20. Northumberland Nye Co., 0.04511.54 Unknown 17115.42 Bonham (1983). (NO) NV

21. Pinson Humboldt Co., 0.1214.11 Unknown 3.513.18 Wallace (in press). (PI) NV

22. Preble Humboldt Co., 0.0913.09 Unknown 1.311.18 Wallace (in press) Antoniuk (PR) NV and Crombie (1982) give 1.5

million tons at 0.064 oz .

Table 8.5--Grade and tonnage data for sediment-hosted, disseminated precious-metal deposits--(continued)

Tonnage Au grade Ag grade (million)

Location troy ozfgram troy ozfgram short tonfmetric ton Notes

23. Rain Elko Co., 0.083f2.85 Unknown 8.317.53 Bonham (1983) Skillings (RA) NV (1984).

24. Relief Canyon Pershing Co., 0.0411.37 Unknown 8.017.26 Bonham (1983). (RE) NV

25. Sanrmy Creek Elko Co., 0.216.86 Unknown 312.72 Wallace (in press). (SM) NV

26. Santa Fe Mineral Co., 0.03f1.03 0.26f8.91 6.3515.76 Anonymous (1984). (SA) NV

27. Standard Pershing Co., 0.04811.65 0.11613.98 0.88410.80 Bonham (1983). (ST) NV

28. Sterling Nye Co., 0.25f8.57 Unknown 0.500f0.45 Bonham (1983). (SE) NV

29. Taylor White Pine Co., Unknown 3.0f102.86 1019.07 Bonham (1983). (TA) NV

30. Tonkin Springs Eureka Co., 0.084f2.88 Unknown 6.5f5.90 Bagby (1983, unpublished (TO) NV field notes).

31. Windfall Eureka Co., 0.034f1.17 Unknown 3.0512.77 Wilson (1976). (WI) NV

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Wilson, W. L., 1976, The Eureka Windfall gold mine, (abs.); j r ~ Geology and Exploration Aspects of Fine-Grained, Carlin-Type Gold Deposits: A symposium presented by The Geological Society of Nevada and MacKay School of Mines, March 26, 1976, University of Nevada, Reno.

Wrucke, C. T., 1985, Gold Acres, Nevada deposit check list; in Tooker, E. W. (ed.), Geologic cha rac te r i s t i c s of t he Sediment- and Volcanic- Hosted Disseminated Gold Deposits--Search for an Occurrence Model: U.S. Geological Survey, Bulletin 1646, p. 120-123.

Wrucke, C. T. and Armbrustrnacher, T. J., 1975, Geochemical and geologic relations of gold and other e lements a t t h e Gold Acres open-pit mine, Lander County, Nevada: U.S. Geological Survey, Professional Paper 860, 27 p.

Chapter 9 RELATIONSHIP OF TRACE-ELEMENT PATTERNS TO ALTERATION AND MORPHOLOGY IN EPITHERMAL PRECIOUS-METAL DEPOSITS

Miles L. Silberman and Byron R. Berger

INTRODUCTION

An epi thermal o r e deposit i s defined a s a relatively near-surface deposit formed in a hydrothermal system under low t o moderate pressure and a t empera tu re range below about 300°C (Barrett , 1985). This concise definition is a r e s t a t emen t of Lindgren's character is t ics of hydrothermal sys tems of "epithermal" character . A modification of Lindgren's character is t ics is tabula ted in Table 9.1. These character is t ics a r e both physical and chemical, and w e will, in this and t h e following paper (Berger and Silberman, 1985, this volume), a t t e m p t t o r e l a t e them.

Epithermal lode deposits in t h e Circum-Pacific region produce approximately 30 million grams of gold annually (Ciles and Nelson, 1982) and a larger, bu t indeterminate , amount of silver. Many epi thermal deposits a r e closely associated with convergent p l a t e boundaries re la ted t o present and relatively r ecen t regimes of p la te t ec ton ic in teract ion (Giles and Nelson, 1982; Sawkins, 1984). These mobile regions of t h e earth's crus t a r e character ized by r ecen t volcanism, high h e a t flow and tec tonic ac t iv i ty , and by t h e presence of ac t ive and recently ac t ive geothermal fields, some of which have deposited precious me ta l s and associated me ta l s (Table 9.1) in similar concentra t ions (but no t volumes) t o those found in t h e epi thermal ore deposits (Weissberg et al., 1979; Henley, 1985, this volume).

The understanding of processes t h a t occur during t h e format ion of epi thermal o re deposits has been advanced in the r ecen t pas t by t h e suggestion t h a t t hese o r e deposits a r e essentially fossil geothermal sys tems (e.g., White, 1955, 1981; White, 1974; Wetlaufer e t al., 1979; Henley and Ellis, 1983; Henley, 1985, this volume). The observed d a t a discussed in t h e references above demonstra te t h a t epi thermal o r e deposits and geothermal sys tems have similar a l ter - ation mineralogy, temperatures , fluid compositions and stable-isotope patterns, and geochemical associations. Indeed, f ea tu res character is t ic of geothermal systems, such a s siliceous s in ter and hvdrothermal explosion breccias a r e found in some epi thermal deposits, and can b e par t of t he o re (Barrett , 1985; Silberman et al., 1979; Wallace, 1980, 1984). The analogy between epi thermal ore deposits and geothermal sys tems is supported by the occurrence of ore-grade concen- t ra t ions of Au and Ag and other associa ted e l emen t s (e.g., As, Sb, Hg, TI, W) in surface discharge mater ia l from several ac t ive geothermal systems, such a s S teamboa t Springs, Nevada, Ohaaki-Broadlands, and Waiotapu, New Zealand, and elsewhere (Weissberg, 1969; Weissberg et al., 1979; Ewers and Keays, 1977;

White, 19811, although t h e amount of th is mater ia l found t o d a t e i s small.

Much published information is available on t h e geology? a l tera t ion mineralogy and zoning, fluid composition and isotopic character is t ics , geometry of hea t sources and flow, and pa t t e rns of fluld migration in geothermal systems. The geology and a l tera t ion mineralogy and zoning, temperature , isotopic composition and salinity d a t a for o re fluids f rom many epi thermal precious-metal deposits a r e also well

Table 9.1--Characteristics that classify a hydrothermal system as being epithermal (after Lindgren, 1933)

Depth of formation Surface to lOOOm

Temperature of formation 50' to 300'~

Form of deposits

Ore textures

Ore elements

Alteration

Common features

Thin to large veins, stockworks, disseminations, replacements

Open-space filling, crustification, colloform banding, comb structure, brecciation

Au, Ag, (As, Sb), Hg, [Te, T1, U], (Pb, Zn, cu)*

Silicification, argillization, sericite, adularia, propylitization

Fine-grained chalcedonic quartz, quartz pseudomorphs after calcite, brecciation

* [ I brackets indicate elements seldom present in more than sub-economic concentrations; (1 parentheses indicate elements often present in economic concentrations but actually less valuable than associated precious metals.

documented (Buchanan, 1981; Heald-Wetlaufer e t al., 1983: Havba e t al.. 1985. t h i s volume). Detailed d a t a on geochemistry and geochemical zoning in both types of sys tems a r e less available, and not nearly a s well documented. The geotherm~l-epi therma! analogy is generally accepted, and mos t models t h a t have been proposed for types of epi thermal deposits draw heavily on t h e l i tera ture on geothermal systems. We will, in th is and the following paper (Berger and Silberman, 1985, this volume), review geochemical zoning in geothermal systems, and present a number of morphological models and descriptions of var ie t ies of epi thermal o re deposits, and summarize our available d a t a on geochemistry, geochemical zoning, and i t s relationships t o a l tera t ion mineralogy and physical morphology.

GEOTHERMAL SYSTEMS

Morphology and Charac te r i s t i c s

White e t al. (1971) def ine a geothermal system a s a source of hea t within t h e ear th ' s crust , be t h a t from a magmat ic intrusion, o r regional hea t flow, and the rocks and water a f f e c t e d by t h a t heat. The geothermal sys tem includes both t h e upwelling hot fluids and t h e marginal convect ive downflowing, cold recharge waters. Stable-isotopic studies of geothermal sys tems have demonstra ted tha t , by far , t h e g rea te s t majority of them contain water predominantly of me teo r i c origin (White, 1974, 1981). The surface expressions of geothermal systems a r e hot springs, fumaroles, and other indications of hydrothermal ac t iv i ty such a s a l t e red rocks. The surface phenomena a r e generally a very small f rac t ion of t h e s ize of t h e geothermal system, perhaps on t h e order of 5 percent (Henley, 1985, this volume).

In order t o gene ra t e a geothermal system, i t is necessary t o have a source of heat , a source of water, and a geologic se t t i ng t h a t provide original or induced zones of permeabili ty t h a t allow the water t o flow and be recharged. The morphology of geothermal sys tems in volcanic areas , such a s a r e common in t h e Circum- Paci f ic region, a r e described in deta i l by Henley and Ellis (1983) and Henley (1985, this volume). Examples of these sys tems include Broadlands and Waiotapu, New Zealand; Steamboat Springs, Nevada; Yellowstone, Montana-Wyoming; and El Tatio, Chile. Henley (1985, this volume) shows typical schemat ic cross sections of geothermal sys tems hosted in silicic volcanic terranes, common on continental margin areas , such a s Yellowstone and Broadlands and andesi t ic stratovolcanic ter ranes , common in island a r c areas, such a s Matsao, Taiwan.

Alteration Pa t t e rns

Alteration pa t t e rns in a var ie ty of geothermal sys tems have been studied in detail. The a l tera t ion mineral assemblage varies with temperature , fluid composition, and host-rock primary mineralogy. Most geothermsl sys tems show la tera l and ver t ica l changes in a l tera t ion mineral assemblages and have considerable variation in thei r spatial distribution.

Since geothermal sys tems evolve through t ime, temporal variations in t h e positions of channels and t h e water t ab le can cause ext remely complex overprinting of a l tera t ion assemblages.

The spat ia l variation of t h e a l tera t ion pat terns a r e i l lustrated by t h e resul ts of studies a t Steamboat Springs, Nevada (White et al., 1964; Sigvaldeson and White, 1961; 1962). The geothermal system occurs in pre-Tertiary metamorphic and grani t ic rocks which a r e overlain by Miocene t rachyandesi te and andesi te flows and pyroclastic rocks, and L a t e Ter t iary and Pleistocene alluvial deposits. These alluvial deposits contain interbedded basal t ic andesi te flows. A ser ies of rhyolite domes in t rude t h e older volcanic rocks and alluvium. The t ime-stratigraphic relations a r e shown in Figure 9.1.

The a l t e ra t ion mineralogy of two drill cores i s shown in Figures 9.2 and 9.3, one (GS-5, Fig. 9.2) collared in sinter in a presently ac t ive par t of t h e system, and t h e o the r (GS-7, Fig. 9.3) collared in an acid-sulfate a l t e red zone in a n a r e a above subsurface boiling. Hole GS-5 probably represents a dominantly hypogene a l tera t ion assemblage (Sigvaldeson and White, 1962). Sinter occurs f rom t h e surface t o 84 f ee t , grading f rom opaline t o chalcedonic with depth. The alluvium which occurs beneath t h e s in ter has been silicified by deposition of hydrothermal qua r t z with illite, K-feldspar, and iron sulfides. The rocks beneath t h e alluvium a r e a l t e red t o an assemblage of K- feldspar, illite, a lb i te , chlorite, and residual quar tz , which gives way at depth, gradually t o a similar assemblage with ca lc i te , g rea t e r amounts of chlor i te and some sericite. In t e rms of hydrothermal mineral

Alluvium (Opalinel

Lake Lahontan alluvium

- Erosional disconformity (Major break ~n the deposition of

slnier and alluvium)

Pre-Lake Lahontan alluvium

< (largely chalcedonic) 7

(Much undeclphered history)

- 9 - C) 0 0 0

Steambosi Hills

Rhvoiite 1.2 m Y. Basali~c andeslte (2.5 m.Y)

Cobbles of

F i g u r e 9.1. Compos i t e t i m e - s t r a t i g r a p h i c rela- t i o n s a t Steamboat Spr ings , Nevada, i n d i c a - t i n g a to ta l l i f e t i m e o f t h e g e o t h e r m a l s y s t e m i n e x c e s s o f 2.5 m i l l i o n y e a r s ( f rom S i l b e r m a n et al., 1979).

assemblages, t h e zoning would correspond t o su r f ace s in ter , underlain by si l icif ied rock, overlying a n argil l ic or phyllic assemblage, which slowly gives way a t depth t o propylit ized rock (Rose and Burt , 1979). In detail , t h e a l tera t ion minera ls occur in irregular proportions, and the re is a gradat ional change in the i r re la t ive proportions below t h e alluvium-volcanic con tac t (Fig. 9.2). Zones of chalcedony and chalcedony-quartz ca l c i t e veining a r e indica ted on Figure 9.2 fo r GS-5. The content of ca l c i t e in t h e veins increases with depth. The q u a r t z ve ins in GS-5 range up t o 8 f e e t in thickness (White e t al., 1964).

In con t r a s t t o GS-5, hole GS-7 (Sigvaldeson and White, 1962) displays a n a l t e r a t ion pa t t e rn cha rac t e r - i s t ic of acid-sulfate leaching near t h e surface (Fig. 9.3). The wa te r t ab l e is a t 114 f e e t in t h e hole, and sub-surface boiling has resul ted in separa t ion of C 0 2 and H2S gas f rom t h e hydrothermal fluid. The H S gas condenses in near-surface wa te r s which a r e oxi&zing, and is conver ted t o sulfur which prec ip i ta tes , and t o sulfuric acid. The resul tant ac id ic water percola tes downwards, and leaches t h e grani t ic rock t o a residual, porous mass of opaline si l ica and residual qua r t z above t h e water table. The acid-sulfate wa te r s a l t e r t h e

grani t ic rock t o a n assemblage of kaolinite, alunite, and q u a r t z below t h e wa te r table. This mineralogy gives way a t dep th t o a montmorillonite-dominated assemblage, with kaolinite, illite, residual feldspar, and qua r t z a s t he ac id is r eac t ed away. With increas ing dep th chlorite, ser ic i te , and i l l i te a r e present. The upper a l te ra t ion zone would be a su l fo tar ic a l t e r a t ion assemblage, succeeded by advanced-argil l ic minerals a t t h e water table , giving way gradually t o argil l ic, and then propylit ic assemblages. As in GS-5, t h e a l tera t ion minera ls occur in variable proportions with depth (Fig. 9.3). A t about 358 f e e t , t h e assemblage would be considered propylitic. A t g r e a t e r depth , t h e assemblage again becornes argillic.

In many geothermal sys tems t h e drill-core d a t a show t h a t a l t e r a t ion i s most in tense (and t empera tu re s highest) around fissures where hot wa te r flows. Miner- alogy grades t o lower t empera tu re assemblages away f rom these fissures. The fissures themselves t end t o conta in quar tz , K-feldspar, and ca l c i t e or wairakite, where boiling has o r i s occurring (Ellis, 1979).

Permeabi l i ty is a n impor tant controll ing f ac to r in a l te ra t ion . Impermeable rocks a r e generally less

Principal minerals, approximate percentage P D ~ oy 2: 4-0 0.1 0.5 1.0 5 10 50 100 c.'> " " Opsiine slnter wlth cristoballte and 1 chaicedony in general increasing 0;-

j[ downward to 80 ft; chalcedonic sinter with cri~tobai#te 80 to 84 ft

D~sconlorm~ty

[Chalcedanlc slnter

-- Unconfarmlty(?) Tuff breccla

- Unconformlty

x xgr x xex e

x xx"x

- Soda trachyte

- dike Granodiorite I Summary geochembstry

Hole GS-5 Steamboat Springs. Nevada

Figure 9.2. Alteration mineralogy and original lithologies in drill hole GS-5, Main Terrace, Steamboat Springs, Nevada. Water table is at or slightly above surface, and springs are still actively flowing. Summary geochemistry for selected elements in sinter and quartz vein from drill core are plotted on the same depth scale (modified from Sigvaldeson and White, 1962).

.I Summary geochem#stry

Hole GS-5 Steamboat Sprmgs. Nevada

6-

Summary geochemistry Hole GS-5 '- Steamboat Sprmgs. Nevada

Principal minerals, * Opal, cristobalite

Opal, amorphous

approximate percentage

T - + O Acid-leached granodiorite

Quartz, relict

Acid-altered granodiorite

Altered granodiorite

Altered pyroxene andesite

dike

Altered granodiorite

P 6

Figure 9.3. Alteration mineralogy and original lithologies in drill hole GS-7, Silica Pit, Steamboat Springs, Nevada. Water table is at -114 feet, and sulfur deposition and solfataric alteration are occurring abve the water table (modified from Sigvaldeson and White, 1962).

a l tered, even in higher t empera tu re zones of geothermal sys tems than more permeable rocks. In hole GS-5 (Fig. 9.2) a t Steamboat , t h e alluvium is completely a l t e red t o an assemblage quartz- Kfeldspar-illite-pyrite, although t h e t empera tu re in

this upper level is only between 130' and 1 4 0 ' ~ (Sigvaldeson and White, 1962).

Geochemical Zoning

High concentra t ions of Au and Ag and associated e l emen t s a r e found in su r face or near-surface deposits in some geothermal sys tems (Ewers and Keays, 1977; Weissberg, 1969; Weissberg e t al., 1979). Table 9.2 l ists chemical d a t a fo r surface discharge mater ia l fo r several geothermal areas. Au and Ag, along with e lements normally associated with them in epi thermal deposits (Table 9.11, a r e strongly concentra ted . In some samples, t he precious-metal values a r e in o re grades, particularly t h e siliceous, antimony-rich precipi ta tes from Steamboat Springs, Nevada, and Broadlands and Waiotapu, New Zealand. However, t h e ac tua l volume of mater ia l enriched t o th is ex ten t is very small in a l l of t h e cases found so far.

Chemical da t a for drill-core mater ia l f rom Steamboat Springs and Broadlands show t h a t t h e group of e lements As, Sb, Hg, B, TI, and Au tend t o be concentra ted in t h e upper pa r t s of t h e geothermal sys tems and decrease with increasing depth. Base meta ls , Cu, Pb, Zn, and o the r e l emen t s Bi, Se, Te, Co, and t o a n ex ten t Ag, appear t o be precipi ta ted a t g rea t e r depth in higher t empera tu re zones. However, t he re a r e deviations f rom this simplified scheme, and these variations a r e significant. F igure 9.2 shows se lected chemical d a t a p lot ted a s a function of depth fo r drill hole GS-5 f rom Steamboat Springs, Nevada, described earlier. The top four samples a r e sinter. The remaining, deeper ones a r e chalcedonic-quartz or chalcedonic quartz-calcite veins cut t ing the a l t e red rocks below t h e sinter. The d a t a show strong enr ichment of Sb, As, and TI in t h e near-surface par t of t h e system. Nearly a l l of t h e TI i s in t h e sinter, with samples below i t having about crus ta l average content for unaltered grani t ic rock. Sb and As both decrease with depth, with Sb concentra ted largely in t h e upper 100 feet . B is strongly enriched near t h e surface; below t h e s in ter i t i s about a t crus ta l average levels fo r grani t ic rock. Hg was no t analyzed in these samples, but Table 9.2, and d a t a discussed by White (1981) indicate s t rong enr ichment of Hg in t h e near surface. No cinnabar is found below 50 f e e t from the present topographic surface. C u and Zn a r e present a t low concentration levels in t h e veins and sinter, generally below crus ta l averages for grani t ic rock. Pb was not de t ec t ed in most samples. Zn shows an irregular increase with depth, C u irregularly decreases with depth. Sr increases irregularly with depth below t h e sinter, and is probably representa t ive of increasing ca l c i t e content of veins with depth (Sigvaldeson and White, 1962; White et al., 1964). Au is de t ec t ab le in some samples of t h e sinter, but i s highest in veins, just beneath i t , then decreases t o below detect ion (0.1 ppm) with g rea te r depth. Ag is also present in t h e sinter, but is highest in veins a t moderate depths, between 84 and 393 f ee t , and appears t o be decreasing again at deeper levels.

There a r e samples of mater ia l precipi ta ted a t depth in Steamboat Springs t h a t have e levated base- me ta l content. Sample 7, Table 9.2, is from a

Table 9.2--Geochemistry of s u r f i c i a l hot-spring d ischarge m a t e r i a l .

F i e l d No. Sample o r

No. d e s c r i p t i o n Area Rock Type Au Ag A s Sb Hg T1 R Ba Re S i Mn Cu Pb Zn Co Se Te Ref

1 BROADLANDS Sb-rich s u l f i d i c - New s i l i c a p p t , 85 500 400 2000 630 NA NA NA NA NA NA 25 70 M NA NA Zealand Ohaki Pool

S i n t e r d ischarge 0.45 13 83 159 ND 3.3 ND ND ND 2.7 ND 20 20 20 ND 0.5 0.09 from wel l

5 F

4 Orange WAIOTAPU Orange ppt 80 175 P, ZJ 170 320 NA NA NA NA NA NA 15 50 NA NA NA PPt New Zealand Champagne Pool

E G

5 Orange PPt

6 S i n t e r S i n t e r (Bulk) 8.4 5.0 NA NA NA NA NA NA NA NA NA NA NA NA NA NA NA 2

Champagne Pool @ m 7 W941C STEAMBOAT Meta-s t ibni te h 60 400 600 2000 80 2000 2000 NA NA NA NA 2000 400 2000 NA NA NA

SPRINGS, o p a l d ischarge P Nevada from 200- w e l l W m

10 MTM

11 S i n t e r H i l l

q S i l i c e o u s mud, 15 50 701) 1.5X 100 700 500 NA NA NA NA 20 7 50 NA NA NA deep l e v e l s

% w

S i n t e r w i t h s t i b n i t e 1.5 1 50 1.OX 30 70 1000 NA NA NA NA 1 NA 0.2 NA NA NA Surface s p r i n g

S u l f i d e mud, s u r f a c e 0.1 1.5 35 340 6.5 54 >ZOO0 300 5 N(10) 100 5 N(10) 10 N(5) NA NA s p r i n g

Chalcedonic s i n t e r , 0.1 N(0.5) N(5) 500 >13 N(.2) 200 300 3 ~ ( 1 0 ) 50 L(5) L(10) L ( 5 ) ~ ( 5 ) NA NA o l d t e r r a c e

[NA = Not analyzed; ND = Not d e t e c t e d ; N(5) = Not d e t e c t e d a t l i m i t i n d i c a t e d i n parentheses ; < = Less than amount i n d i c a t e d ; > - Greater than amount i n d i c a t e d ; v a l u e s i n ppm u n l e s s o therwise indica ted . ]

l ~ e i s s b e r g , 1969; ' ~ w e r s and Keys, 1977; 3 ~ e d e n q u i s t , 1984; 4White, 1981; 5 t h i s r e p o r t . Analysts: John Gray, James Domenico.

discharge blas t f rom a 650-ft (200-111) well, consisting of opal and meta-s t ibni te and has quite significant Cu, Pb, and Zn content . Sample 8, which has moderate amounts of base metals, is believed t o be a precipi ta te from deep levels carr ied physically t o t h e surface during maximum discharge. Both of these samples are , in f ac t , a lso very high in Au, Ag, Sb, As, and Hg-- e lements normally found nea r t h e surface. These two samples f rom deep within t h e system a r e t h e only mater ia l approaching precious-metal ore grade a t Steamboat Springs. Unfortunately, t he geometry and s ize of any possible zone of mineralization t h a t they come from cannot b e determined f rom t h e available data . However, if a deep zone of mineralization exists a t Steamboat Springs, i t does not appear t o conform t o t h e t race-metal zoning indicated by the unmineralized drill co re GS-5, but appears t o b e strongly enriched in both base me ta l s and precious meta ls and e lements considered cha rac te r i s t i c of deposits of t h e la t ter .

EPITHERMAL ORE DEPOSITS

Morphology and Charac te r i s t i c s

Epi thermal o r e deposits occur in a wide variety of types and geologic se t t ings within the broad, global f ramework mentioned earlier. Classification schemes have been based upon p la t e t ec ton ic concepts (Giles and Nelson, 1982; Sillitoe, 1981; Bonham and Giles, 1981; Sawkins, 19841, association with volcanic landforms (Henley and Ellis, 1983; Sillitoe and Bonham, 1984)) physical and mineralogical character is t ics (Lindgren, 1933; Giles and Nelson, 1983; Heald- Wetlaufer et al., 1983; Hayba e t al., 1985, this volume), and associa ted magma types and mineralogy (Bonham, 1986). Berger and Eimon (1983) proposed conceptual models of volcanic-hosted epi thermal systems based on a combination of physical, mineralogical, and hydrological characterist ics. Some s tudents of epi thermal deposits f i t a l l var ie t ies of them into a gener ic model, flexible enough t o f i t a l l t he variations (Buchanan, 1981; Silberman, 1982). We discuss the morphology of epi thermal sys tems using a hybrid scheme of classification based largely on physical and mineralogical (observational) characteristics, following Berger and Eimon (19831, but modified from recent work of Giles and Nelson (1982); Bethke (1984); Heald-Wetlaufer et al. (1983); Bonham (1986); and Ashley and Berger (1985, in press).

Epithermal ore deposits may occur in a continuum of types ranging from shallow quartz-pyrite stockworks and breccias--the hot-spring environment-- t o relatively deep veins and fissures, t he bonanza environment. I t is very important t o emphasize the possible continuum of f ea tu res and character is t ics because models proposed for these deposits usually represent end-member cases, and neglect t o portray the very important in termedia te cases. Table 9.3, modified f rom t h e discussion of Giles and Nelson (19821, Bonham (19861, and Ashley and Berger (1985, in press), presents t h e important character is t ics of t h e deposit types.

In addition t o sources of hea t and fluids (as in geothermal systems), t h e precious-metal epi thermal

deposits need a source of metals, a mechanism of m e t a l t ranspor ta t ion t o a place of deposition, a mechanism of precipitation, and a sufficient t i m e for t h e t ranspor ta t ion and deposition mechanisms t o ope ra t e so t h a t economic concentrations of t he meta ls can occur. This question of t iming is a cr i t ica l one, and will be discussed in a la ter sect ion following the brief description of a f ew epi thermal sys tems and thei r geochemical and a l tera t ion variations. The na tu re of t h e me ta l sources and mechanisms of transportation and deposition a r e covered elsewhere in th is volume (Henley, 1985; Henley and Brown, 1985; s e e also Henley and Ellis, 1983; Browne and Ellis, 1970).

Deposit models for these o r e sys tems have been proposed by many authors (cf. Buchanan, 1981; Henley and Ellis, 1983; Hayba et al., 1985, this volume; Bonham, 1986). The ear ly representations tended t o be qui te general, and were designed t o f i t a l l of t h e observed pa t t e rns of mineralization and a l tera t ion and me ta l associations in one context. Late ly (Berger and Eimon, 1983; Giles and Nelson, 1982; Bonham, 1986; Ashley and Berger, 1985, in press), t h e diversity of types of epi thermal deposits has been emphasized, and most authors of t h e recent papers propose multiple models fo r t h e general epi thermal class of deposits. Schemat ic drawings of t h e deposit-type models a r e generally presented a s cross sections through the idealized systems, and show more or less deta i l on t h e pa t t e rns of a l tera t ion and various physical f ea tu res observed in t h e deposits. Most of t h e authors accep t t h e geothermal-epithermal analogy, and connect t h e various deposits t o surficial f ea tu res of geothermal sys tems a t t h e t o p of t h e cross section. All of t hese cross sect ions suffer from thei r two-dimensional orientation, and do not adequately portray the th ree dimensional character is t ics , and variations t h a t a r e present in t h e ac tua l sys tems (see, fo r contras t , Vikre, 1985).

Deposit-type models a r e useful fo r focusing a t t en t ion on changes of mineral assemblages and physical character is t ics of t h e deposits with depth, and form a f ramework within which t o discuss geochemical variations. I t must be kept in mind, however, t h a t in individual deposits, a s in t h e geothermal systems, genet ic processes t a k e place in an ext remely complex manner, in which t h e e f f e c t s of ver t ica l and la tera l variations in t empera tu re due t o such e f f e c t s a s positions of t he water table, boiling zones, and channels of fluid movement occur. The resul tant mineral a l tera t ion and me ta l deposition pa t t e rns a r e made even m o r e complex by changes in these e f f e c t s with t ime.

Figures 9.4, 9.5, and 9.6 present ra ther generalized models from Berger and Eimon (1983) and Ashley and Berger (1985, in press) for a var ie ty of bonanza and hot-spring-type volcanic-hosted precious- me ta l deposits. Disseminated replacement or "Car1in"-type sys tems a r e discussed by Bagby and Berger (1985, th is volume).

Al tera t ion Pa t t e rns

Study of t h e a l tera t ion mineralogy and zoning in epi thermal deposits is progressing rapidly, but t h e amount of deta i led information is not y e t a s available

M. L. SILBERMAN & B. R. BERGER

Barren atz+calc~te*f luori te*barl te

Figure 9.4. Schematic cross section of quartz-adularia or low sulfur bonanza deposit, Bonanza- IA model, showing alteration mineralogy and two zones of mineralization from the "closed cell convection" model of Ber- ger and Eimon (1983).

Figure 9.5. Schematic cross section of the quartz-alunite, or high sulfur bonanza deposit, Bonan- za-IB model, showing the alteration mineralogy and possible location of ore bodies, and other features from Ashley and Berger (1985, in press).

Qfz+illite+native metals+sulfarsenides+sulfantirnonides+ Ag-sulfides+base-metal sulfides

Bl B o n a n z a a n d / o r s t o c k w o r k o r e s

\ \ '+~tz+adulariat i l l i te+~g-sulf ldes+base-metal sulfides

base-metal sulfldes+adularla

\+vein s t r u c t u r e

Base o f pervasive a c i and si l ic i f icat ion

Argillic alteration Ore body Quartz, kaolfnlte, rnantrnorllion~te, Enargkte, luzonlte, tetrahedrlre, mlxed layer clays, pyrlte tennantite, covell~te, bismuthln!t old, tellurides

Advanced arg~ l l ic alterat ion Quartz, alun~te, pyrophyllhte, d~aspore. kaollnlte, pyrite

Propyl i t ic a l tera t ion

Y Propyl i t lc alterat ion

Propyl i t ic a l tera t ion

-Quartz-sulfide veins with K-feldspar alterat ion Fe. Cu. Mo sullldes

210 CHAPTER 9

Table 9.3--Characteristics of epithermal precious-metal systems

Deposit Examples with grade, Type Characteristics tonnage, and/or production 1

Bonanza A) Quartz-adularia or low sulfur type (Fig. 9.4) Vein and Lode Vein and hanging wall stockworks

of quartz, calcite, and adularia. Pyrite is major sulfide, with variable amounts of silver sulfides, sulfosalts, native Au, and electrum. Host rocks are usually andesite dacite, rhyolite. Associated elements are As, Sb, variable Hg, variable Cu, Pb, Zn, Mn. Alteration patterns usually silicification, with K-spar proximal, to argillic then propylitic distal, but geometry varies.

Al) Comstock Lode, NV

(Ag-rich quartz vein and hanging tall s tockwork) 8.3 x 10 oz Au 200 x lo6 oz Ag from 19 million tons

A2) Pachuca, Mexico

(quartz-vein system, base-

k 6 metal ric ) 6.2 x 10 oz Au 1500 x 10 oz Ag from 100 million tons 3.5% Cu, Pb, Zn

A3) Creede, CO

(Ag-bearing base-metal vein system in southern art of 8 district) 0.15 x 10 oz Au 85.2 x lo6 oz Ag with approximately 7.5% Pb, Zn, Cu from 5 million tons

B) Quartz-alunite or high sulfur type (Fig. 9.5)

A variant of this type, the Goldfield B1) Goldfield, NV or enargite-pyrite gold deposits, occur as tabular lodes or pipes of silicified (quartz-alunite breccia breccia and some veins. Ore is native reefs) 4 2 x lo6 oz Au Au or electrum associated with Cu b 1.7 x 10 oz Ag from sulfosalts, and sulfides, pyrite, and 5 1/2 million tons complex sulfosalts and tellurides. Host rocks are usually andesite and dacite. B2) Sumitville, CO Associated elements are Cu, As, Ag, Ba, Bi, Zn, Pb, Te, variable Hg. Alteration (quartz-a unite pods) is advanced argillic, with alunite in i 0.24 x 10 oz Au lodes, zoning outwards to argillic 0.34 x lo6 oz Ag halos, then distal propylitization. 0.427 x lo6 lbs Cu

Both systems structurally focused by major faults, locally controlled by intersections and inflections.

l~uchanan, 1981; P. Bethke, USGS, personal communication, 1985, cf. Hayba et al., 1985, this volume.

in the literature as for the porphyry copper and molybdenum deposits. The general patterns of alteration around bonanza- and hot-spring-type systems are shown in the models (Figs. 9.4, 9.5, and 9.6). An outer propylitic zone is nearly ubiquitous, but the nature and mineralogy of the argillic-phyllic zones is highly variable from system to system and depends on temperatures and fluid compositions (particularly

pH, and a +). In order to document some of the variations $at occur, Table 9.4 describes the physical nature of several deposits, along with the mineralogy of the alteration zones, from proximal to the veins or ore zones to distal.

The variability in the alteration zoning in the bonanza and hot-spring systems can result in difficulty when trying to use alteration to determine depth or

Table 9.3--Characterisitcs of epithermal precious-metal systems--(continued)

Deposit Type

Hot Springs or silicified quartz stockworks

Di s- seminated replacement

Characteristics

A) Quartz-adularia or low sulfur type (~ig. 9.6)

Silicified host rock, breccias, and stockworks of quartz-sulfide mineralization. Native gold and electrum of micron size associated with pyrite, marcasite, and silver sulfosalts. Sulfide content varies. Host rocks are andesite, dacite, or rhyolite. Host rocks usually affected by silicification with adularia and/or albite, variable calcite, or dolomite. This grades outwards to argillic and/or zeolitic alteration, then to propylitic assemblages. Mineralization due to repeated, episodic, and explosive stockwork veining and brecciation (hydrothermal). Associated elements: Hg, T1, As, Sb, Ba, (W). Cu, Pb, and Zn usually deep--all highly variable. Sinter frequently present, and may be mineralized.

B) Quartz-alunite or high sulfur type

These deposits are similar to the A) type except that alteration assemblages are advanced argillic (alunite-bearing) central, zoning out to argillic peripheral zones, and Cu sulfosalts are present in the ore-sulfide assemblages.

A) Tabular, fault-related stratabound orebodies with micron-sized native Au on pyrite and/or organic carbon particles. Host rocks are thin-bedded calcareous clastic sedimentary rocks--limestones, dolomites, or limy siltstones. Variable silicification and carbon remobilization. Variable amounts of brecciation and veining. Jasperoid development along faults which control fluid access and/or along bedding as replacements, above, below, or within ore. Associated elements: Hg, Tl, As, Sb, Ba (w).

Examples with grade, tonnage, and/or production 1

Al) Round Mountain, NV2

195 x lo6 tons at grade of 0.043 opt, Au (reported productiog to 1982-- 0.85 x 10 oz Au)

A2) Borealis, N V ~

6 2.3 X 10 tons at grades of 0.10 opt, Au; 0.5 opt, Ag

(~ot-spring system, hosted in ophiolitic rocks o 6 Mesozoic age) 22 x 10 tons at a grade of 0.16 opt, Au

B1) Paradise Peak NV3

10 x lo6 tons at a grade of: 0.35 opt, Au; 4.7 opt, Ag

Al) Carlin, NV4

11 x lo6 tons of 0.i9 opt, Au (produced) 6.2 x 10 tons of 0.20 opt Au (reserves)

A2) Pinson, NV4

4.2 x lo6 tons of 0.18 opt, g Au (produced) 3.3 x 10 tons of 0.12 opt Au (reserves)

A3) Windfall, NV4

3 x lo6 tons of 0.035 opt Au

' ~ i n ~ l e ~ and Berger, 1985. 3 ~ . Nelson, consultant, written communication, 1985. 4 ~ . Miller, consultant, written communication, 1982; Bagby and Berger, 1985, this volume

proximity t o zones of mineralization. Indeed t h e which gives way a t greater depth to illite-K mica- argillic or phyllic a l tera t ion zones a re very complex. qua r t z assemblages. Adularia is usually present in t h e Changes in geometry and (or) positioning of hea t low-sulfur sys tems in these zones and is a very sources with t ime, sealing of permeable zones by common const i tuent of t h e silicified zones in t h e low- deposition of silica, calcite, e tc . will cause overlapping sulfur hot-spring systems, where i t also tends to occur of a l tera t ion assemblages. In general, upper level in t h e stockwork veins. Solfa tar ic hypogene a l tera t ion assemblages have illite-montmorillonite, with quar tz , and supergene a l tera t ion can overprint ear l ier pat terns

Table 9.4--Alteration mineral zoning in epithermal precious-metal deposits

Deposit Type Character of deposit Alteration mineral zoning

Coms tock Quartz, quartz-pyrite quartz- Sericite-quartz adjacent to stockworks, Lode, NV adularia stockworks in hanging wall zoning to chlorite-illite albite f (Bonanza of major fault. Minor sulfides adularia, accompanied occasionally by vein, IA) (pyrite), sulfosalts, and electrum. strong silicification. Outer zone is

Host rock is andesite. propylitic, albite-epidote-chlorite- calcite-pyrite-zeolites. Near-surface, proximal advanced argillic assemblages (quartz-alunite) zone outward to k a o l i n i t e - p y r o p h y l l i t e - d i a s p o r e - q u a r t z , sericite-quartz, illite-quartz, montmorillonite-quartz, and propylitic.

Aurora, NV' Steeply dipping, massive, fine- (Bonanza grained, multiple-banded quartz- vein, IA) adularia veins. Barren, coarse-

grained quartz cuts early fine- grained productive quartz with replaced lamellar calcite. Electrum with pyrite ore minerals. Veins commonly bounded by stockwork zone in brecciated andesite host rock. Veins are along faults.

Deeper levels--Adjacent to veins q u a r t z - a l b i t e - a d u l a r i a - i l l i t e grades outwards to quartz-albite-adularia- chlorite*calcitefpyrite. Shallower levels--Wide zone in hanging wall of quartz-illite-montmorillonitef adulariafkaolinite. Footwall assemblage grades from quartz-illitefadularia to q u a r t z - i l l i t e - a d u l a r i a - m o n t m o r i l l o n i t e f chlorite to an outer envelope of albite- quartz-illite-chlorite-montmorillonite- calcite. Propylitic assemblage developed regionally.

Oatman, A Z ~ Large, tabular, multi-stage Silicification immediately above ore (Bonanza quartz-calcite-adularia veins, shoots-quartz flooding and quartz- vein, IA) strongly banded. Very low sulfide calcite stockworking. Illite, minor

content, ore is submicroscopic montmorillonite in hanging wall, less wide electrum. Host rocks are andesites in footwall. This zones out to propylitic and rhyolites. alteration, consisting of chlorite-pyrite

assemblages. Illitic zone specifically characteristic of ore shoots.

Gold- Quartz-alunite reefs, consisting field, N V ~ of brecciated and fractured, ( Bonanza altered dacite, along faults. vein, IB) Ore is Au with sulfides and complex

sulfosalts. Occurs generally interstitial to breccia fragments and as coatings. Mineralization occurs only in the "silicified" zones.

Round Mt. , NV (Hot-Spring IIA)

Steeply dipping quartz and quartz-adularia veins;

, lamellar quartz after calcite. Low-angle quartz and quartz- breccia veins; silicified breccias. Native gold and electrum with pyrite; realgar. Disseminated ore in permeable beds. Host rock is

Silicified zones consist of quartz- alunite-kaolinite in varying proportions. Pyrophyllite locally present and increases in proportion with depth. This assemblage grades to quartz-kaolinite-illite, then to mont.-kaolinte-quartz, then to a pervasive zone of propylitic alteration, developed regionally. Limonite present in all zones, after pyrite.

Upper zone of pervasive silicification overlying zone of quartz-illite-chlorite alteration. Veins of adularia crosscut both of above but are most abundant in latter. Deeper zone of quartz-chlorite. Lateral zone of quartz-chlorite-white mica-calicte alteration. Blocks of sinter in central breccia pipe indicate previous

rhyolite ash-flow tuff. presence of sinter overlying system.

Tab le 9 .4 - -Al t e r a t ion m i n e r a l zoning i n e p i t h e r m a l p rec ious -me ta l d e p o s i t s - - ( c o n t i n u e d )

Depos i t Type Charac t e r of d e p o s i t A l t e r a t i o n m i n e r a l zon ing

Hasbrou k S i l i c f i e d and a d u l a r i z e d b r e c c i a s Chalcedonic s i n t e r i n t e r b e d d e d w i t h M i . , NV and s t e e p l y d ipp ing q u a r t z s i l i c i f i e d h o s t rocks . C e n t r a l (Hot-Spr ing, and q u a r t z - a d u l a r i a v e i n s . funnel -shaped zone o f q u a r t z - a d u l a r i a IIA) N a t i v e gold and p y r i t e , a l t e r a t i o n g r a d i n g outward and downward

a c a n t h i t h e i n s i l v e r - r i c h v e i n s . i n t o q u a r t z - a d u l a r i a - i l l i t e . Deepes t a l t e r a t i o n i s q u a r t z - a d u l a r i a - a l b i t e - i l l i t e . L a t e r a l t o q u a r t z - a d u l a r i a - i l l i t e i s q u a r t z - i l l i t e and t h e n q u a r t z - i l l i t e - rnon tmor i l l on i t e .

l ~ u d s o n , 1984. ' s t one and Osborne , 1984; Osborne and S t o n e , 1985. 3 ~ u r n i n g and Buchanan, 1984. 4 ~ s h l e y , 1974. 5 ~ r a n e y , 1985

Hydrothermal explosion breccia

Figure 9.6. Schematic cross section of the quartz-adularia, or low- sulfur hot-springs deposit, or Hot Springs-IIA model showing alteration mineralogy, generalized geochemical associations, and other structural and mineralogical features, from Ber- ger and Eimon (1983). The high-sulfur analog of this type of system has central advanced argillic + alunite assemblages, surrounded by argillic and propylitic halos, and is termed Hot-spring-IIB model in text.

. A S Au. AQ. Hg In seams

Opalized rock or porous, vuggy Natlve S, c lnnabar

Pervasive silicification Dispersed AS, S b Au, AQ, TI

Stockwork velns Au. Ag. As, TI. Sb In qtz, chalcedony

Kaol ln l te, a lun~ te , sll lca, l a ros l t e

Breccia dikes Hydrothermal brecciation [ low-angle velnsl Au, Ag, As, Sb, TI s u l l ~ d e s and q t z

of a l tera t ion , part icularly in t h e upper levels of t h e systems. One consequence of t hese complications i s t h a t t hey seriously a f f e c t geochemical pa t t e rns a s t h e acid-leaching tends t o remove many of t h e pathfinder e l emen t s t o o r e and disperses t hem laterally, or down t o deeper levels, depending on t h e local hydrology.

The Bonanza-IA model (Fig. 9.4) a t t e m p t s t o dep ic t a two-stage development fo r zones of mineralization. The competing e f f e c t s of ver t ica l and l a t e r a l changes in tempera ture , due t o t h e var ia t ions ment ioned above, and t o migration of main-fluid movement channels with t i m e a s old channels sea l and new f r ac tu re s open, due t o t ec ton ic or even solution

Silica

A", AQ, As, (Cu, Pb. Z n l In sulfides wl lh adular la

-0tz-sulfide veins CU. Pb. Zn. IAu, Ag l In sul f ldes wl lh c h l o r ~ t e

processes, can cause addit ional complications in t h e development of ac tua l a l te ra t ion assemblages and the i r distr ibution in a l l of t h e deposit types.

NATURE O F FLUIDS INVOLVED IN GEOTHERMAL SYSTEMS AND EPITHERMAL ORE DEPOSITS

Stable-isotopic studies, par t icular ly isotopic r a t i o s of oxygen and hydrogen in geo the rma l sys t ems and epi thermal o re deposits have shown t h a t t h e hydrothermal fluids in both environments a r e predominantly me teo r i c wa te r in origin (e.g., Taylor,

1973, 1974; White, 1974; Bethke and Rye, 1979; OfNeil and Silberman, 1974). Fluid-inclusion studies in these ore deposits suggest a t empera tu re range of f rom 150' t o 3 0 0 ' ~ for deposition. The fluids in most of t h e sys tems a r e dilute, with NaCIe of 0.5 t o 5 wt. percent. The high-sulfur or qu&tz-alunite sys tems tend t o have somewhat higher salinities, up to about 10 t o 12 wt. percent, and t empera tu res near t h e upper end of the scale (Hayba, 1983; Bruha and Noble, 1983).

The chemistry of these ore-forming solutions and t h e na tu re of t he physico-chemical conditions of me ta l t ranspor t and deposition in both epi thermal o r e deposits and geothermal sys tems is t r ea t ed at length elsewhere in this volume (Henley, 1985, this volume; Henley and Brown, 1985, this volume). Hayba e t al. (1985, th is volume) describe t h e na tu re of t h e geochemical environment associated with the Creede and Summitvil le epi thermal sys tems in Colorado.

TIMING

Mention has been made of variations of temperature , position of zones of fluid movement, and level of water table and zones of boiling with t ime. Time is an ext remely important f ac to r in the development of epi thermal o r e deposits, and evidence for t h e duration of hydrothermal processes and even t s is contradictory. Geochronological da t a , largely K-Ar ages on volcanic sequences hosting mineralization and on t h e products of hydrothermal a l tera t ion of wall rocks, and hydrothermal deposition of gangue minerals in veins, breccias, et ce t e ra , summarized by Silberrnan (1983) and Noble and Silberman (1984) suggest t h a t hydrothermal ac t iv i ty in epitherrnal o re deposits, geothermal systems, and porphyry copper deposits las ts between about one-half t o g rea te r than t w o and one-half million years with an ave rage on the order of 1.25 m.y. In addition, Silberrnan (1983) suggested t h a t most epi thermal deposits can be r e l a t ed in t i m e to e i ther a s t age of local volcanic ac t iv i ty or t o a regional pat tern of distribution of volcanic rocks, which can be in terpre ted within t h e t i m e f ramework of p la te t ec ton ic evolution.

On t h e other hand, from solution composition theory, Henley (1985, this volume) e s t ima tes t h a t quite significant amounts of precious me ta l s can be deposited from even dilute fluids (such a s those in the low-sulfur epithermal systems) on t i m e scales of t he order of lo4 years, although th is is s t a t ed a s a minimum figure. These t i m e es t ima tes appear t o be in conflict with t h e evidence f rom t h e geochronological studies and from detailed studies a t Creede, Colorado, by Barton et al. (1977). However, t h e geochronological studies completed t o d a t e t r e a t a mining dis t r ic t a s a whole, and do not focus on putting t i m e l imi ts on t h e duration of any particular s t age o r episode of a l tera t ion or me ta l deposition. The re is evidence from the physical f ea tu res of epi thermal deposits t h a t mineralizing events a r e short-lived, but appear t o be repetit ive. Silberman (1983) c i t ed severa l lines of evidence for t h e occurrence of episodic mineralizing events in bonanza vein and hot-springs systems: I. S tages of brecciation and stockwork veining can

be shown t o be multiple f rom cross-cutting

relationships, and no t a l l s t ages a r e associated with o re deposition.

2. Epithermal qua r t z veins a r e frequently repetatively banded, and only some bands carry sulfides and precious metals.

3. In many deposits, a l tera t ion assemblages overlap or succeed e a c h o the r in relatively res t r ic ted areas, and only some of those a l tera t ion episodes appear t o b e re la ted t o episodes of o r e deposition.

Many epi thermal o r e deposits a r e spatially associated with small s tocks on t h e order of 2 km * radius, where t h e h e a t of this crystall izing magma is inferred t o drive t h e convect ing hydrothermal system (White, 1974, 1981; Norton and Cathles, 1979). Based on hea t loss modeling studies, Norton and Cathles (1979) indicate t h a t hydrothermal cells driven by this s ize of h e a t source will decay well within about a 100,000-year t i m e span. The geochronological studies a t a number of epi thermal vein deposits and hot-spring deposits suggest l i fe spans of one-half t o one and one- half m.y. (Silberman, 1983; Noble and Silberrnan, 1984). Given the aforement ioned geochronological evidence, additional sources of h e a t must be supplied t o the sys tems beyond a single smal l stock or dome t o keep t h e sys tems active. We suggest t h a t t h e magmat ic sys tems responsible for driving the hydrothermal cells must b e much larger than those surface-exposed s tocks of 2 km t radius, and t h a t t he ac t iv i ty must be pulsed and renewed a t intervals. Evidence for this pulsing comes from the geochronology of Julcani, Peru (Noble and Silberman, 1984), where e ight s t ages of volcanic and interspersed hydrothermal alteration-mineralization episodes occurred over a to t a l duration of 700,000 years, with individual s tages taking between 100,000 and 200,000 years. We conclude t h a t mineralizing episodes may, in f ac t , be short-lived, but occur within a much longer framework of volcanic evolution and hydrothermal activity.

GEOCHEMICAL ZONING IN EPITHERMAL DEPOSITS

Any epi thermal o r e deposit viewed a t t h e present resembles a s t i l l photograph in an ins tant of geologic t i m e t h a t ca t ches a very dynamic, evolving, three- dimensional interplay of hea t sources, fluid flow, chemical reactions, and s t ructura l changes which have combined t o produce a n economic concentration of me ta l s t h a t will be fu r the r modified with additional t ime by post-hypogene processes. Measurements we make of the geochemistry of a n epitherrnal o r e deposit r e f l ec t this snapshot view, and the observed complexities of t h e e l emen t distribution pat terns caution against t h e in terpre ta t ion of the da ta in a "cook book" fashion. We f e e l t h a t i t is unlikely t h a t a single, all-inclusive e l emen t distribution model will be derived t h a t will allow a quick and accura t e determinat ion t h a t a n economically viable mineral deposit lies within any geochemical or favorable lithological anomaly. A t best , we will probably define a reas of re la t ive favorabili ty for fur ther investigation.

Geochemical zoning in epi thermal systems can be studied on a var ie ty of scales, including regional, district , and orebody or ore-shoot dimensions. Ideally,

M. L. SILBERMAN & B. R. BERGER 215

pat terns of geochemical zoning, once they a r e established, should be usable guides t o location of additional orebodies. Most models of epithermal deposits suggest generalized geochemical zoning pat terns (Figs. 9.4, 9.5, and 9.6) t h a t presumably may be applied t o a l l epi thermal o re deposits. However, we believe t h a t t hese zoning pa t t e rns a r e specific t o individual mineralizing systems, and t h a t a unique, widely applicable model of e lement zoning t o delineate ore in a l l districts will probably not be achieved. We have found t h a t t h e r e i s considerable variation in the pathfinder e lements associa ted with precious-metal mineralization, particularly in thei r la tera l and ver t ica l changes in concentration. For example, a s shown in the following section, t h e generalization of decreasing Hg concentra t ion in a l t e red rocks and veins with depth does no t always hold.

In any particular deposit or district , the geochemical zoning pat terns do tend t o have regularity, and once established, can serve a s guides t o additional reserves or s i tes of mineralization in t h a t district. Tha t is, they have application t o prospecting only a f t e r thei r relationships t o known mineralization a r e established. We will i l lus t ra te zoning pat terns a t a variety of d is t r ic t a n d orebody scales f rom actual examples of deposit geochemical studies, in t h e following section of th is report , and in Berger and Silberman (1985, this volume) t o follow.

BODIE MINING DISTRICT

The Bodie mining district , Mono County, California, produced approximately 1.5 million ounces of Au from banded quartz-adularia veins t h a t c u t Miocene andesites, daci tes , and pyroclastic rocks in one of several Ter t iary volcanic cen te r s in t h e region (Kleinhampl e t al., 1975). The mining dis t r ic t is an erupt ive center whose major s t ruc tu re i s an irregular, faulted, north-trending anticline formed by intrusion and doming of t h e flows and pyroclastic rocks by small dac i t e intrusions (Fig. 9.7). The intrusive rocks occupy vents from which t h e volcanic rocks were erupted. Several s e t s of s teeply dipping f au l t s c u t all of t h e rocks exposed, including t h e intrusions. One prominent s e t s t r ikes N t o NE, and another i s normal t o this. The major veins and f r ac tu res a t Bodie also s t r ike N t o NE within one of t h e f au l t sets. Most of t h e production c a m e from the northern par t of t he district , in the vicinity of a small graben filled with tuff breccia t h a t was faul ted down into t h e intrusive dac i t e of Bodie Bluff during or shortly a f t e r i t s emplacement (Fig. 9.7).

The productive quar tz veins c u t both the extrusive and intrusive rocks. The veins vary in thickness from about I t o 90 f e e t , although most a r e not more than a f ew f e e t thick. The o re minerals a r e principally native gold and silver. Argentite, pyrite, and sphalerite a r e present and increase in abundance with depth a s t h e tenor of t h e gold decreases (Chesterman e t al., in press). In t h e main productive zone, old records quoted by Chesterman et al. (in press) indicate t h a t t h e o r e averaged 1.7 ounces per ton Au, and 3.1 ounces per ton Ag. Adularia is a common const i tuent of t h e mineralized quar tz veins,

with K con ten t s in t h e range of approximately 3 t o 9 percent. The adular ia somet imes fo rms euhedral crys ta ls up t o 3 c m long along open f rac tures (Silberman e t al., 1972; Silberman, 1983; O1Neil et al., 1973).

In the southern pa r t of t he district , and reportedly a t deep levels in t h e main "bonanza" zone in t h e southern par t of t he Bodie Bluff area , a vein system rich in base me ta l s and silver i s reported (Chesterman et al., in press). Mineralogy of this vein system is more complex than t h e simpler quartz- adularia veins t o the nor th and south and consisted of an assemblage of sulfides and sulfosalts. Dump samples of this ma te r i a l assay up t o 2 percent Cu and Pb, 1.4 percent Zn, and 220 ppm Ag (M. L. Silberman, unpublished data , 1980).

Wall-rock a l tera t ion assemblages in t h e northern pa r t of t h e Bodie Bluff a r e a a r e being determined by petrographic and x-ray di f f rac t ion analyses by Pe te r A. Herrera of The Colorado School of Mines. Figure 9.8 is a s chemat i c cross sect ion along line XX1 of Figure 9.7, summariz ing preliminary results of t hese determinations f rom surface outcrops, underground samples, and drill c o r e supplied by t h e Homestake Mining Co. An upper zone of quartz-adularia a l tera t ion of t h e intrusive and extrusive rocks gives way a t greater depth t o a n assemblage of adularia- illite-quartz. This deeper a l tera t ion assemblage is character is t ic of wall-rock mineralogy in the main bonanza production zone. A t even g rea te r depths, and peripheral t o the zone of economic mineralization, propylitic assemblages occur. The graben a t Bodie Bluff contains silicified fall-out and fall-back breccias of hydrothermal origin, which contain a variety of c las ts of intrusive and extrusive rock, and siliceous sinter. Argillic a l tera t ion occurs beneath t h e silicified breccias (P. A. Herrera , wri t ten communication, 1985). Hematite-quartz-adularia matr ix hydrothermal breccias with K-silicate a l t e red c las ts c u t t h e quartz- adularia a l t e red rocks a t t h e top of Bodie Bluff. The breccias and thei r host rocks a r e themselves c u t by chalcedonic, banded qua r t z veins, which a r e mineralized (Silberman, 1982, 1984; P. A. Herrera , wri t ten communication, 1985).

Silberman (1982, 1984) suggested t h a t t he Bodie mining dis t r ic t is a mineralized geothermal center with most of i t s upper near-surface f ea tu res preserved. The a l tera t ion pat terns a r e qui te similar t o those of t h e Hot Springs A model, and Herrera 's studies, including t h e important documentation of t h e presence of s in ter in explosion breccias a t Bodie Bluff, appear t o confirm this suggestion. In all of t h e previous s tudies a t Bodie, t he sinter was not identified. Bodie has f ea tu res character is t ic of both Hot Springs A and Bonanza A deposits. I t was mined largely fo r bonanza-type lode qua r t z veins, but t he re a r e numerous s e t s of these, and large, structurally focused vein sys tems such a s those a t Oatman or t h e Comstock Lode a r e not present. A broad, blanket-like zone of quartz-adularia a l tera t ion, accompanied by hydrothermal breccias and mineralized stockwork veining a r e cha rac te r i s t i c s frequently found in t h e hot- spring environment.

Detailed stable-isotopic analyses of t he mineralized qua r t z veins and a l t e red wall rocks f rom

-Contact - - Fault

'8 Alluvium In Q) 0)' .- I - a >; a i? .Q (3 C

a3

E % sg - B 5

Andesite Q Q

= E 0 i $ 0 .- z EL $ z E : : 2 $ Tuff breccia 2 2

:2 a 0 a0 r 0

Figure 9.7. Generalized geology and cross California (modified from Silberman et

Intrusive dacites :: 3 a0

cn V)- '

Extrusive dacites .gz Dacite flows and intrusives

h'o , Tuff breccia (II,! $ Tuff breccia

~m .;; 0) e

I

sections of the Bodie mining district, Mono County, al., 1972).


Schematic alteration 'cross-section, Northern Bodie Bluff

Stock work quartz veining

Silicified fallout breccias, bedded fallback breccias, 90 Fe oxides, relict pyrite, quartz - contains sinter

ic alteration; kaolinite, montmorillonite,

Sericitic alteration

8 8

r fJY 0 z % n

86

84

Modified from P A Herrera, unpublished data, 1985 - Fault Veins are not shown in actual positions, and

Vein their distribution is hypothetical C _ / Alteration contact

Figure 9.8 Schematic alteration cross section of the northern part of Bodie Bluff (section XX') showing alteration mineralogy and relationships between important structural and hydrothermal features. Samples from Zone 111, the Upper Hobart Tunnel, and the W i e Tunnel form the basis for the geochemical discussion in the text (modified from P. A. Herrera, USGS, unpublished data, 1985).

t he K-silicate and propylit ic zones and fluid-inclusion t empera tu re and salinity determinat ions (O'Neil e t al., 1973; Nash, 1972) indicate t h a t mineralization and a l tera t ion were produced by fluids of me teo r i c origin t h a t were hea ted t o t empera tu res of between 215' and 245OC, with salinit ies of l e s s than 0.5 %-equivalent NaCI. These o r e fluids a r e isotopically and chemically similar t o present-day geothermal fluids in t h e Bodie Hills a r e a (O'Neil et al., 1973).

OINeil et al. (1973) discussed major and se lected minor e l emen t chemical relations of a l tera t ion in the Bodie Bluff area; however, they did not a t t e m p t t o delineate ver t ica l o r l a t e ra l zoning patterns. They documented t h a t K-silicate a l tera t ion, which accompanied emplacemen t of t he ore-zone veins, resulted in a n e t gain in K 2 0 , a loss in Na20, CaO, and MgO, and l i t t l e change in S i 0 2 and A120 . Sr was depleted and R b gained during the process. f he degree of a l tera t ion appeared to O'Neil et al. (1973) t o co r re l a t e with depletion of ''0, and was explained a s progressive replacement a n d recrystall ization of t h e host rocks by K-feldspar (adularia) of similar chemical and isotopic composition as t h a t of t he veins. P. A. Herrera is current ly complet ing a deta i led t race- e lement and major-element chemical study of the Bodie Bluff a r e a t h a t will a t t e m p t t o r e l a t e a l tera t ion assemblages and e l emen ta l chemical variation in detail.

Trace-element analyses a t Bodie were carr ied out by OINeil et al. (1973) during the isotopic and major-element variation sampling. The results of some of these analyses were briefly discussed by Silberman (1982, 1984). Da ta from th ree s e t s of samples col lec ted over a ver t ica l range of approximately 600 f e e t a r e summarized in Figures 9.9 and 9.10. Q u a r t z veins and a l t e red wall rock were sampled from the quartz-adularia a l tera t ion zone a t t h e top of Bodie Bluff (zone 111 of Fig. 9.8) a t elevations of 8,900 t o 9,000 feet ; in the K-silicate (adularia-illite-quartz) zone of t h e Upper Hobart Tunnel at a n elevation of 8,740 feet ; and, in the propylitic a l tera t ion zone of the Bodie Tunnel a t an elevation of 8,400 f ee t , a s s t ructura l ly deep within the d is t r ic t a s i s current ly accessible. The Upper Hobar t Tunnel is in t h e productive zone, and one vein (approximately 2 f e e t wide) in the presently accessible par t of t he workings was stoped. The Bodie Tunnel was largely a haulage adi t , but a moderately large vein (approximately I foot wide) was followed with a dr i f t and partly stoped. Quar t z veins in the Upper Hobart Tunnel are , with the exception of the one stoped, on t h e order of a f ew inches in width. The veins in t h e Bodie Tunnel a re , in general, t h e s a m e width t o even thinner. Quar t z veins in zone 111 a r e normally on t h e order of one t o th ree inches in width (occasionally wider) and a r e vuggy, porcellaneous, and micro- crystall ine t o chalcedonic in texture . The deeper veins a r e f ine grained, vuggy, and sheeted. All veins have considerable l imonite along f r ac tu res and vugs.

The d a t a fo r s e l ec t ed t r a c e e lements in vein samples and wall rock f rom zone 111 and t h e Upper Hobart and Bodie tunnels a r e plotted a s bar graphs on a logarithmic scale (Figs. 9.9 and 9.10). For t h e Upper Hobart and Bodie tunnels, t he plots a r e organized from t h e center (neares t t h e cen te r of Bodie Bluff) t o t h e

periphery of t h e mineralized zone, f rom top t o bottom of t h e figure. The adi ts a r e projected on to t h e line of section, a s both a r e slightly north of it.

Figures9.11 and 9.12 show trace-element contents of samples collected at se lected in tervals la tera l ly away f rom stoped, mineralized quar tz veins in t h e Upper Hobart Tunnel and Bodie Tunnel. The plots a r e organized with t h e e lementa l concentration of the vein in t h e cen te r and t h a t of the surrounding wall rock on e i the r side. The zoning for Bodie Tunnel (BT) and Upper Hobart (UH) a r e p lot ted together in Figure 9.12 s o t h a t d i f ferences in e lementa l content a s a function of depth (and a l tera t ion assembla e can be 58 . compared. Figure 9.1 1 also shows t h e 6 0 wlth dis tance from t h e vein in the upper Hobart tunnel. The t w o s e t s of plots (Figs. 9.9, 9.10, 9.11, and 9.12) allow a comparison of geochemical zoning t o be made on both a d is t r ic t and ore-shoot scale--a scale d i f ference of, in this case, about one order of magnitude.

We show t h e trace-element d a t a for descriptive purposes of delineating e lementa l zoning in an epi thermal sys tem and make no conclusion here about t h e physico-chemical causes of this zoning. The ac tua l a l tera t ion mineral assemblage of the samples from e a c h sui te a r e shown on Figure 9.8, t h e schemat i c cross section of Bodie Bluff.

Large-scale Vertical Zoning a t Bodie Bluff-- t h e Big P ic tu re

Quar t z veins--The general pat tern of e lementa l concentra t ion changes a r e well i l lustrated by t h e bar graphs (Figs. 9.9 and 9.10). Gold is present in chalcedonic qua r t z veins in zone 111 in significant quantit ies, usually between I and 10 ppm. Gold is highest in t h e qua r t z veins of t h e Upper Hobar t Tunnel, and decreases in concentra t ion in veins of t he Bodie Tunnel level t o generally less than a ppm, although the majority of veins sampled have concentra t ions above t h e level of determination. Figure 9.9 shows t h a t t h e gold content of t he veins is higher towards t h e periphery of t h e d is t r ic t than in veins towards t h e cen te r of Bodie Bluff. Silver is highest in t h e qua r t z veins of zone 111, where i t reaches concentra t ions of between 10 and 100 ppm. I t drops slightly in t h e veins of t h e Upper Hobart , and is lower s t i l l in the Bodie Tunnel, although a few high concentra t ions in samples a r e found there. Silver also tends t o be higher in t h e peripheral par t of t he Bodie Tunnel, re la t ive t o t h a t par t of t he tunnel nearer t h e c e n t e r of Bodie Bluff. Arsenic tends t o be slightly higher and more consistent in t h e quar tz veins of t h e Upper Hobar t re la t ive t o those of zone 111, but is highest in t h e qua r t z veins of the Bodie Tunnel. Antimony, although more irregular in concentration, essentially follows t h e s a m e pattern. Mercury is very variable, reaching i t s highest single sample concentra t ions in the Bodie Tunnel, but i s basically much t h e s a m e level of concentration in general in a l l t h r e e levels.

The base meta ls , Cu, Pb, and Zn, a r e a t low levels of concentra t ion in t h e qua r t z veins a t Bodie, and a r e a t o r near levels of crus ta l average fo r i n t e rmed ia t e rocks (Parker, 1967). Copper is about t h e

same concentra t ion in t h e qua r t z veins of zone 111 and t h e Upper Hobart , and i s highest in t h e interior pa r t of t he Bodie Tunnel. Lead increases from zone 111 veins t o t h e Upper Hobart , and again, i s highest in t h e interior pa r t of the Bodie Tunnel. Zinc is present in consistently measurable amounts only in t h e Bodie Tunnel. The pat tern of Mn concentration is complex. I t is most consistent i n t h e veins of zone 111, although individual samples a r e higher in the Upper Hobart. I t reaches i t s highest levels of concentration in t h e peripheral zone of t h e Bodie Tunnel, whereas i t tends t o be lower in general in t h e interior par t of t h e Bodie Tunnel t han in the t w o shallower zones. Barium is about t h e s a m e in veins of zone 111 and the Upper Hobart , and is highest in the Bodie Tunnel veins. Strontium is present consistently only in t h e veins of t h e Bodie Tunnel. I t i s generally below detectabi l i ty in t h e Upper Hobart , and present a t low levels in zone I11 veins.

In summary, in t h e qua r t z veins of Bodie Bluff, Au is highest a t i n t e rmed ia t e depth levels, although significant concentra t ions occur in the shallow level, zone 111. Silver dec reases with depth, and is on average lower in t h e productive zone veins than nearer t he paleo-surface. Arsenic and Sb increase with depth in the system; Hg shows l i t t l e variation in concentration. Copper, Pb, and Zn increase with depth, a s do Ba and Sr. Manganese distribution is complex, and highest in the deep, peripheral zone of the Bodie Bluff system. There also appears t o be a s t rong la tera l zoning in t h e deeper, propylit ic pa r t of t h e system with C u and Pb highest near t h e cen te r of t he bluff, and Au, Mn, and Ba ( the l a t t e r weakly) higher in t h e peripheral par t of th is a l tera t ion zone. Other e lements , such a s B, Cr , and Ni were not de t ec t ab le in enough samples for pa t t e rns t o be established.

I t is ins t ruct ive t o compare t h e variation pat terns of t h e Bodie qua r t z veins t o the silicious deposits of Steamboat Springs, Nevada (Fig. 9.21, including s in ter and qua r t z veins. The Bodie pat terns a r e different. In particular, t h e e lements As, Sb, and Ag show a pa t t e rn of ver t ica l zoning di f ferent from t h a t in t h e S teamboa t Springs geothermal system, whereas Sr, Cu, and Zn, which increase with depth a t Bodie, a r e similar t o t h e zoning observed a t Steamboat Springs. Mercury, which is strongly concentra ted near t h e surface at Steamboat , maintains about t h e same level of concentra t ion throughout the Bodie Bluff hydrothermal system.

Al tered wall rocks--The pat terns of ver t ica l variation in e lementa l concentration of t h e wall rocks a t Bodie a r e highly irregular and t rends a r e more difficult t o identify (Fig. 9.10). The d a t a for zone 111 a r e separa ted in to a s e t of results for samples collected adjacent t o qua r t z veins, or containing quar tz s t r ingers (upper bar graphs) and those not adjacent t o qua r t z veins (lower bar graphs). Gold and Ag in zone 111 rocks a r e highest in samples adjacent t o or containing qua r t z veins, whereas As and Sb a r e highest in rocks away f rom quar tz veins. I t is r a r e for Au t o be g rea te r than 1 ppm and Ag g rea te r than 10 ppm in wall rocks anywhere in the Bodie Bluff par t of t h e system, although a f e w higher concentrations do occur. Cold contents a r e quite irregular, and a r e highest in t h e wall rocks of zone 111, in t h e rocks

adjacent t o o r containing qua r t z veins. Gold is s t i l l present in significant levels in t h e Upper Hobar t Tunnel, and drops off considerably a t t h e Bodie Tunnel level, al though most samples st i l l have de tec t ab le amounts present. Thus, gold appears t o irregularly decrease with depth in t h e system. The highest Ag values a r e found in wall rocks adjacent t o veins in zone 111 and in t h e wall rocks of t h e Upper Hobart Tunnel, and a r e lower elsewhere. Mercury is irregularly high in both t h e Upper Hobar t and Bodie tunnels, and lower in zone 111. C u is consistently present, although irregular, a t about t h e in termedia te composition rock crus ta l ave rage in zone 111 and Upper Hobart Tunnel wall rocks, and lower in t h e Bodie Tunnel. Zinc and Mn appear highest in t h e Bodie Tunnel wall rocks. The other e l emen t s a r e just t oo irregular in concentration t o specify variation.

Thallium was not analyzed in the OINeil et al. (1973) sample suite. P. A. Herrera (writ ten communication, 1985) repor ts TI a t highest concentra t ions in a l t e red wall rocks, particularly si l icified-hematit ic breccias, in zone 111 (with a range of concentra t ion of about 5 t o 15 pprn), with concentra t ions generally decreasing a t greater depth in t h e system.

Detailed La te ra l Zoning

Ore-shoot scale variation in e lement concentra t ions a t distances on t h e order of inches and f e e t f rom mineralized qua r t z veins was studied by two sampling t raverses , approximately perpendicular t o two veins--one vein, two f e e t wide, which dips a t about 85' SW, in t h e Upper Hobart Tunnel, and another, about one foo t wide, which dips 60°SW in t h e Bodie Tunnel. Several o ther sampling t raverses around veins in o ther accessible underground workings a r e in t h e process of being analyzed and compiled (P. A. Herrera , unpublished data , 1985).

Figure 9.1 1 f rom d a t a of OINeil et al. (1973) shows t h a t t h e r e is some r e g a r i ty in the distribution Yh of K, Rb, Sr, and Au and 6 0 rela t ive t o t h e two- foot-wide, mineralized qua r t z vein in t h e Upper Hobar t Tunnel. A channel sample of t h e vein contained 4.2 ppm Au. The wall rocks a r e considerably lower than t h a t a t t h e vein margins, and drop off laterally. The e l emen t s R b and K, which increase during K-silicate a l tera t ion, increase a s t h e veip is approached, and Sr decreases. In like manner, 6 80 becomes more negative a s t h e vein is approached representing depletion of t h e heavy isotope of oxygen. Figure 9.12 shows summary plots of two longer t raverses around t h e s a m e Upper Hobart vein, and a vein in t h e Bodie Tunnel, t h a t were dr i f ted along and stoped in places. These plots do no t show t h e regularity of t e la tera l variation shown by the K, Rb,

f 8 Sr, Au, and 6 0 data. In most of these samples, t h e gold con ten t was below t h e .05 ppm detectabi l i ty l imi t and was no t plotted. Silver falls off rapidly in t h e wall rocks with t h e exception of a high-grade sample in t h e hanging wall of t he Bodie Tunnel vein. This sample had qua r t z veinlets, and is mineralized with 1 5 ppm Au a n d e levated contents of Cu, Zn, As, and Hg a s well. The distribution pat terns of most of t he e l emen t s plotted a r e too irregular t o serve a s guides t o t h e

Bodie quartz veins

Upper Hopart p I

4- Bodie T u ~ n e l Y 8400

Figure 9.9. Selected trace-element content of quartz veins from three different vertical levels in the northern part of Bodie Bluff.

proximity of mineralized veins with t h e exception of Hg, and possibly Mn. Manganese fo rms a halo around t h e Bodie Tunnel vein t h a t is be t t e r developed in t h e footwall from about 5 t o 20 f e e t from t h e vein. Mercury forms a pronounced halo a t about 10 t o 25 f e e t from both veins ( the l imit of t h e traverse).

The profile Sam ling a t t he deta i led sca l e (inches t o f ee t ) shows t h a t l s 0 depletion, K and R b addition, and a decrease in Sr indicate proximity t o a

mineralized vein, and t h a t t h e veins appea r t o have a Hg halo surrounding them. Manganese appears t o halo t h e vein in t h e propylitized zone, but no t t h e vein in the K-silicate a l t e red productive zone. On the district-wide sca l e of ver t ica l zoning, i t would be very difficult t o predict one's proximity t o t h e zone of economic mineralization from t h e t race-e lement data--particularly t h e pathfinder e l emen t s t h a t t he typical epi thermal models suggest can b e used for this

Bodie quartz veins

ppm

1 Upper H?bart 8740

I I ?--- Bodie Tu,nnel 8400

Figure 9.9. ( a n t 'd. )

purpose. In f ac t , in t h e quar tz veins a t Bodie, t h e pathfinder e lements As and Sb tend t o increase with depth, and ra ther than being concentra ted a t shallow levels above t h e zone of mineralization, Hg does not vary much with depth. In t h e wall rocks, t he con t ra s t s in these e lementa l concentrations with depth, if present, a r e very weak. On t h e sca l e of t h e a r e a of Bodie Bluff (less than 20% of t h e a rea of t h e mining district) , i t appears t h a t a Mn halo might be t h e mos t

useful geochemical indicator of a direction towards economic mineralization. T o recognize th is halo, qua r t z veins should be sampled and analyzed and not whole-rock samples.

Although the trace-element pa t t e rns a r e quite complex, a pa t t e rn of zoning appears t o be present a t Bodie such t h a t a s cheme of qual i ta t ive evaluation could b e s e t up. Good signals fo r t h e approach t o the zone of economic mineralization would be increasing

Bodie wall rocks ppm

Zone Ill 8900'-9000

Upper Hobart --u Bodie Tunnel

8400'

EZ2 Adjacent to or with quartz veining 0 No quartz veining

Figure 9.10. Selected trace-element concentrations for altered wall rocks from three different vertical levels in the northern part of Bodie Bluff. For Zone 111, samples adjacent to or containing quartz veins are plotted separately from those with no veining. Alteration assemblages of the wall rocks are given on the schematic cross section of northern Bodie Bluff (Fig. 9.8).

Bodie wall rocks

ppm

Upper Hqbart 1 8740

Bodie Tunnel 8400'

Adjacent to or with quartz veining 0 No quartz veining

Figure 9.10. (cont'd.)

Sr. ppm 0

BOO

600

400

200

0

Variation of Au. K. Rb. Sr. 8180 with distance from vein sample UH 3 A - Upper Hobart Tunnel

E W

Au and slight increases in As, Sb, and Hg in qua r t z veins. This should be associated with a sl ight dec rease in Ag, and slight increases in P b and Ba. Evidently, when Sr and Zn a r e consistently present, and Pb, Ba, and Mn a r e a t maximum values, t h e o re zone is above t h e a r e a being sampled (Figs. 9.9 and 9.10). These exploration geochemical guides, however, may be useful only because they have already been empirically re la ted t o t h e ac tual zone of productive mineralization. One could argue t h e point t h a t t he bes t indicator of economic Au mineralization a t Bodie i s Au in t h e qua r t z veins!

Geochemical zoning pat terns a t Bodie do not conform, in to ta l , t o those predicted by t h e epi thermal models o r t h e geothermal analogy, but t hey do have some regularity, and could with fu tu re ref inement be t e s t ed a s predictive tools.

The a l tera t ion mineral assemblages and physical f ea tu res of t h e rocks appear to be much b e t t e r guides t o ver t ica l and la tera l position in t h e district . Upward widening of t h e zones of hematite-quartz-adularia matr ix breccias a r e useful for delineating near-surface supra-vein positions. These breccias a r e much like similar f ea tu res in zones of si l icification above o re shoots a t Oatman, Arizona (Durning and Buchanan, 1984). A t Bodie, although they cer ta in ly occur above t h e general bonanza vein zone, we have not y e t been ab le t o associa te one of these breccias with a specific, subsurface vein.

Figure 9.11. Variation of the concentration of Au, K, Rb, Sr, and oxygen isotope ratio with distance from a two-foot-wide quartz vein in the Upper Hobart Tunnel, northern W i e Bluff.

A Channel of quartz vein

PARAMOUNT MINING DISTRICT-- VERTICAL ZONING

The gold prospect in t h e Paramount mining district , located about 5 miles NNE of Bodie, is hosted in rhyolit ic pyroclastic and epic las t ic rocks and dac i t e lava flows. The a rea produced Hg (Kleinhampl e t al., 1975). A t t h e surface, extensive zones of silicification, brecciation, and argil l ic a l tera t ion a r e exposed. Siliceous sinter occurs a s large disaggregated blocks and is exposed near t h e old Paramount Hg workings. The Homestake Mining Company and Houston Oil and Minerals have drilled t h e prospect fo r t h e possibility of a disseminated hot-spring-type volcanic-hosted gold deposit. The Homestake Mining Company allowed us t o sample and analyze a 425-foot reverse circulation ro t a ry drill hole t h a t encountered a zone of Au mineralization about 50 f e e t thick averaging 1.4 ppm Au. The hole was collared in alluvium near t h e blocks of sinter, and passed in to volcanic breccias and pyroclastic rocks t o a depth of 230 feet . Below 230 f e e t i t encountered dac i t e and andesi te flows, and remained in t h e rocks t o to t a l depth. The alluvium and volcanic rocks a r e a l t e red t o assemblages character- i s t ic of moderate t o in tense silicification, argil l ic a l tera t ion, and chlorit ization (Homestake Mining Co., wr i t ten communication, 1985). We have not yet done x-ray petrographic mineral assemblage determinations of t h e drill cut t ings and have only generalized d a t a for

M. L. SILBERMAN & B. R. BERGER 225

a l tera t ion assemblage determination. The drill-hole logs indicate t h a t severa l zones of chalcedonic quar tz veining a r e present, and a n in ter face between a zone of oxides (hemat i te and jarosite) and a zone of sulfides plus oxides (pyrite plus hemat i t e and l imonite) occurs a t about 170 f e e t depth. Sulfides, principally pyrite, vary from about 10% immediate ly beneath the oxide- sulfide in t e r f ace t o about 2%, generally decreasing with depth (Fig. 9.13).

Figure 9.14 summarizes available chemical da t a for 5-foot in tervals of t h e hole a s bar graphs.

The zone of gold mineralization occurs between 110 and 160 f e e t depth, where t h e Au content is consistently above I ppm, with a maximum value of 1.9ppm, immediately above t h e oxide-sulfide interface. P a r t of t he gold-mineralized zone contains 5 t o 10 percent chalcedonic quar tz veins, but t h e res t of t he zone has no chalcedonic veins l isted on the drill log. Alteration in t h e zone of gold mineralization is also variable, consisting of a l ternat ing zones of moderate t o in tense silicification and argillization. Deeper in the hole (215-240 feet) , a 25-foot zone of Au mineralization varies f rom about 0.5 t o I ppm, and is coincident with the occurrence of I t o 5 percent chalcedonic quar tz veining in silicified rock. A few other in tervals of 5 or 10 f e e t with g rea te r than 0.5 ppm Au intervals occur lower in t h e hole, but these intervals appear unrelated t o e i ther sulfide content o r qua r t z veining. Although irregular, higher concentrations of Au occur beneath the "mineralized" zone than above t h e zone near t h e surface, where i t increases irregularly towards t h e zone.

Silver is generally low near t h e surface in t h e drill hole, although some high intervals a r e found, and shows a very irregular pa t t e rn throughout t h e res t of the hole. I t appears t o be slightly lower in t h e gold- mineralized zone (a t about 2 ppm) than t h e 50-foot interval in t h e hanging wall and the 75-foot interval in t h e footwall, where t h e Ag contents vary from 3 t o 5 ppm. There is a g rea t e r consistency in t h e Ag concentration in the 7 5 f e e t or s o below t h e oxide- sulfide in ter face where about 10 percent sulfides a r e present. Near t h e bottom of t h e hole, a f ew intervals have a s high a s 5 t o 20 pprn Ag and a r e unrelated t o any obvious veining or a l tera t ion type.

The mercury concentra t ion is high in most of t he hole, a s i t is elsewhere in t h e Paramount d is t r ic t (Kleinhampl e t al., 1975; M. L. Silberman, unpublished data , 1985). The Hg content in t h e drill hole is above 10 pprn near t h e surface decreasing t o about 50-foot depth, where i t increases again t o g rea te r than 10 pprn. The higher values ( > I 0 ppm) appear t o occur in t h e 50-foot interval above t h e gold-mineralized zone. I t is s t i l l a t o r near 10 pprn for about 100 f e e t below t h e mineralized zone and then irregularly decreases t o to ta l depth. Within t h e Au zone p e r se, t h e Hg content is lower (4 t o 8 ppm) than in t h e hanging wall of t h e zone, and slightly lower than the proximal par t of t he footwall. The Hg distribution suggests a Hg halo t h a t is be t t e r defined in t h e hanging wall.

Thallium concentrations a r e generally high, two t o t en t imes t h e crus ta l average fo r felsic rocks (Parker, 1967). Thallium is lowest in concentration and irregular in distribution near surface, and increases a t a depth of 50 f e e t and remains mostly

above 10 pprn f o r another 50 feet . I t decreases again in the Au-mineralized zone, then increases strongly just below t h e zone where t h e concentra t ion remains mostly above 10 ppm fo r approximately 100 feet . This footwall zone of high TI concentration corresponds approximately t o sulfide con ten t s above 10 percent. The thallium increases again near the bottom of the hole, but with t h e except ion of one 5-foot interval, is below 10 ppm. Thallium appears t o halo t h e Au zone, with a sharper peak t o t h e halo in t h e hanging wall.

Arsenic and ant imony both have complex pat terns of ver t ica l distribution and a r e consistently above 200 and 100 ppm, respectively. Arsenic tends t o decrease from near t h e surface t o t h e Au-mineralized zone, where i t varies between 500 and 1,000 ppm. I t falls sharply a t t h e oxide-sulfide in ter face and increases again in t h e zone of 10 percent sulfide beneath the in ter face . Below the high-sulfide zone i t decreases, and increases again near t h e bottom. Antimony is distributed in a similar pat tern , but shows a more general dec rease with depth, in ter rupted by t h e mineralized zone where a slight positive anomaly occurs in pa r t of t h e zone. Low concentrations in the footwall As and Sb appear more re la ted t o the oxide- sulfide in t e r f ace than t o t h e gold mineralized zone.

The boron concentra t ion is 100 pprn just below t h e surface, and decreases irregularly with depth. The gold-mineralized zone in ter rupts the decrease and has a poorly defined halo of relatively high B surrounding it. Barium is slightly lower in the Au-mineralized zone (700 t o 1,500 ppm) than above and below it. The highest concentra t ions (2,000 t o 3,000 ppm) a r e found in t h e 35-foot in terval just above t h e Au-mineralized zone. Strontium also appears t o be a t a minimum concentration in t h e Au-mineralized zone, and increases in both hanging and footwalls with the highest concentra t ions occurring in t h e footwall, increasing irregularly t o about 100 f e e t above TD, then decreasing again. The Sr pa t t e rn is reminiscent of t h a t found for l a t e ra l variation of Sr around t h e vein in t h e Upper Hobar t Tunnel a t Bodie (Fig. 9.1 1).

Molybdenum is highest in the Au-mineralized zone reaching levels of 30 t o 50 ppm. I t is generally a t lower concentrations both above and below t h e zone. Outside of t h e Au in terval t h e highest Mo contents a r e where t h e sulfide con ten t i s highest. Molybdenum is relatively high a t Paramount , whereas a t Bodie and many of the o ther epi thermal sys tems from which we have data , i t i s present at very low levels, although a t Round Mountain, Aurora, and Divide, Nevada, Mo is present in relatively significant concentrations.

Manganese forms a well-defined halo in the hanging wall of t h e gold-mineralized zone where i t reaches values up t o 2,000 pprn in t h e 60-foot interval above the zone. I t is a t a minimum within t h e zone and increases for a 60- t o 70-foot interval below i t with concentrations between 1,000 and 5,000 ppm. Manganese drops off below this level, although a n occasional high interval is encountered.

Copper and lead concentra t ions have relatively similar distribution patterns. They increase irregularly from the surface t o about 100-foot depth, a r e relatively low in t h e Au-mineralized zone, and increase below this, particularly in the interval with t h e highest sulfide content. Both C u and Pb form an

Upper Hobart 8740' A Bodie Tunnel 8400'

Upper Hobart 8740' A Bodie Tunnel 8400

I I+~oundary of vein I I 10 10 20 Feet 10 10 20 Feet

Foot wall Hanging wall Foot wall Hanging wail

Figure 9.12. Variation of the concentration of selected trace elements with distance from a mineralized quartz vein in the Upper Hobart Tunnel, and in the Bodie Tunnel, northern Bodie Bluff. Hg at both levels, and Mn in the Bcdie Tunnel appear to halo the vein.

ostensible halo around t h e Au-mineralized zone, but largely because low values a r e found within th is zone. Zinc irregularly decreases f rom near su r f ace t o t o t a l depth , excep t f o r irregular highs, which break t h e pa t t e rn in t h e Au-mineralized zone. C o i s near t h e de t ec t ion l imi t a t t h e surface , f i r s t is d e t e c t a b l e a t a dep th of 60 f e e t , and then decreases towards t h e Au- minera l ized zone, where i t is below detection. Below t h e zone, C o i s a t relatively high values where t h e sul f ide con ten t i s high, and then decreases with depth. T h e n e a r t o t a l absence of C o in t h e Au-mineralized zone makes i t appear t o form a halo around t h e zone, bu t t h e C o distribution is too irregular t o be useful a s a pathfinder t o gold mineralization. Nickel has i t s highest concentra t ion within and just below t h e zone of highest su l f ide content and shows l i t t l e relationship

t o t h e Au-mineralized zone. Chromium also shows l i t t l e relationship t o t h e Au-mineralized zone, and has i t s highest concentra t ions midway down t h e hole, in a l a rge in terval (150 f e e t ) t h a t includes both t h e Au and high-sulfide zones. Vanadium dec reases irregularly f rom near t h e su r f ace t o about 100 f e e t above t o t a l depth , where i t increases slightly again.

In summary, t h e distribution of some e l emen t s (Au, As, and Sb) appea r s t o be r e l a t ed t o t h e oxide- su l f ide in ter face , while o the r distr ibutions (Ag, Cu, Pb, Co, and Ni) appear t o be re la ted t o t h e t o t a l sulfide content , and o the r s (Hg, TI, Mn, B, Sr, and Ba) appea r t o form halos of varying degrees of definit ion around t h e Au-mineralized zone. Vanadium and z inc show l i t t l e definable relationship t o e i t he r f ea tu re , t h e sulfide o r gold-mineralized zones.




Mn

+Boundary of vein

Cu

+Boundary of vein

20 Feet Fool wall Hanging wail Foot wall Hanging wall

Figure 9.12. (conc'd.)

From t h e perspective of an explorationist , proximity t o t h e gold-mineralized zone is best indicated by Hg, TI, and Mn and t o a lesser e x t e n t by Ba and B. In th is hydrothermal system, proximal increases in Mn, TI, and Hg, and perhaps Ba, below a B-enriched zone might serve a s exploration geochemical guides t o gold mineralization. We feel t h a t o ther halos, such a s those for Pb, Cu, Co, and Sr a r e too poorly defined t o b e valuable in exploration. We a r e perhaps most intrigued by the in ter rupt ion of t h e t race-e lement distribution pat terns by t h e partly chalcedonic-veined, Au-mineralized zone. The in terpre ta t ion of the t race-e lement pa t t e rns is compl icated by many factors--the oxide-sulfide in ter face , t h e high-sulfide zone, and t h e presence of qua r t z veining. A t Paramount, we have not a s ye t

r e l a t ed t h e e lement distribution pa t t e rns t o a l tera t ion mineralogy, and our understanding and in terpre ta t ion of t h e pat terns will perhaps change when this p a r t of our research is completed.

SUMMARY

The study of geochemical zoning pa t t e rns in epi- t he rma l o re deposits i s important t o evaluat ing t h e use of geochemical pathfinders f o r locating potent ia l ore- grade mineralization. T o accomplish this objective, i t is important t o be able t o r e l a t e t h e physical and geochemical morphologies of t h e hydrothermal systems. Conceptual models of epi thermal ore-deposit types a r e of necessity generalizations of t h e geometry,

Paramount, California DDH PRC 82-3

Sulfide content, %

" O Volcanic clasticsl '. volcanic pyroclastics

:- Andesite dacite

Alteration Silicification

I Intense Weak-moderate

Argillization lntense Weak-moderate

1 Chloritization ,#.* .... Quartz veining, %

F i g u r e 9.13. G e n e r a l i z e d lithology, a l t e r a t i o n minera logy, s u l f i d e c o n t e n t , and p e r c e n t a g e o f c h a l c e d o n i c q u a r t z v e i n i n g i n dr i l l h o l e PRC 82-3, f r o m the P a r a m o u n t prospect, W i e H i l l s , Mono County, C a l i f o r n i a ( m d i - f i e d f rom i n f o r m a t i o n s u p p l i e d by Homestake M i n i n g Co.).

a l tera t ion, s t ructura l controls, o re and gangue mineralogy, and trace-element geochemistry of these deposit types. Variations in t h e pat terns due t o duration, s t ructura l complexity, e f f e c t s of host-rock properties such a s permeabili ty and chemical reactivity, t he overprinting of f ea tu res by changes in the position of ac t ive vents, e levat ion changes in t h e wa te r table, and changes in t h e levels of boiling make t h e ac tua l relationships in deposits very complex. How complex t h e ac tual pa t t e rns can be is i l lustrated by chemical variations, l a t e ra l and ver t ica l a t Bodie, California, and ver t ica l a t Paramount , California. In addition, t he geochemical pa t t e rns in ac tual deposits can be qui te d i f ferent from t h a t reported in ac t ive geothermal systems. From t h e geochemical results at Steamboat Springs, Nevada, and Bodie and Paramount, California, we conclude t h a t t h e presence of mineralization in these epi thermal systems causes deviations from an all-inclusive, s tandard "geothermal zoning" of elements. The deviations in t h e t race- e lement pat terns a r e re la ted t o t h e physico-chemical processes t h a t occur during t h e deposition of t h e ores,

and if t h e processes c a n be demonstra ted t o be regular, perhaps t h e t race-e lement pa t t e rns will be useful for predicting t h e location of ore-grade mineralization in any given epi thermal system.

Both t h e Bodie and Paramount hydrothermal sys tems i l lus t ra te ver t ica l trace-element zoning in. only a portion of a mineralizing system, and, a t t ha t , on t h e scale of hundreds of fee t . The lateral-zoning example a t Bodie i l lus t ra tes t h e types of pat terns t h a t can b e expec ted around individual o r e shoots. In order t o derive geochemical pat terns a s guides to mineralization at t h e district-wide scale , much more d a t a than were summarized above a r e necessary.

In t h e following paper, Berger and Silberman (1985, this volume) summarize t h e results of detailed chemical analyses of ro t a ry and diamond drill holes from two hot-spring-type gold deposits, Round Mountain and Hasbrouck Mountain, Nevada. In both studies, drill holes which t ransected t h e mineralized zones and thei r weakly mineralized-altered margins were analyzed chemically and the trace-element results were combined with d a t a on a l tera t ion mineralogy. The combination of mineralogy and t race- e l emen t chemis t ry produces a be t t e r large-scale, three-dimensional representa t ion of the physical and chemical morphologies of hydrothermal sys tems than we were able t o present f rom t h e essentially two- dimensional study a t Bodie, and t h e one-dimensional study a t Paramount. Round Mountain and Hasbrouck Mountain also show considerable variation from a standard "geothermal system" me ta l zoning model, but both also demons t r a t e regularity in thei r e lementa l distribution pat terns t h a t might be of predictive significance.

ACKNOWLEDGMENTS

We would like t o express our appreciation t o Pe te r Herrera of t h e Colorado School of Mines for providing us his information and insights on the geology of t h e Bodie Mining District , and t o Bob Blakestad of Homestake Mining Company for providing access and information on Bodie and on the Paramount prospect. Analytical d a t a summarized within this report were provided over a 17 year period of t i m e by many chemis ts of t h e Branch of Exploration Geo- chemistry of t h e U.S.G.S. We a r e particularly indebted t o R. M. O'Leary for providing we t chemistry (acid digestion-AA), and M. S. Erickson for providing Emission Spectrographic d a t a for t h e many hundreds of samples submit ted fo r analyses. The manuscript has benef i t ted markedly f rom cr i t ica l reviews by Bob Blakestad and Andy Wallace, but we bear sole responsibility fo r t h e conclusions reached within.


Downhole geochem~stry, Paramount. California

lo3 I

B

E 0

8 --- ern ro a, sg 82

p%%{ sulfides o 2

lo-' 1 10

-

Oxides - - - - - - - - -0x1des 8.

sulfides

-

-

*High grade Au zone. 1.4 ppm, USGS

Figure 9.14. Summary geochemistry of selected trace elements of 5-foot intervals of cuttings from drill hole PRC 82-3, from the Paramount prospect, Mono County, California, plotted as a function of depth. DL is detection limit for the analytical method used. 1/2 DL is one half that value.

Oxides

3-

4 - &

Downhole geochemistry, Paramount, Cal~fornia

ppm 1

Oxides - - - - - - - - - Oxides & sulfides

*High grade Au zone. 1.4 ppm, USGS

Figure 9.14. (cont'd.)

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M. L. SILBERMAN & B. R. BERGER 23 1

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232 CHAPTER 9

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Sillitoe, R. H., 1981, O r e deposits in Cordil leran and island-arc sett ings; & Dickinson, W. R., and Payne, W. D. (eds.), Rela t ions of Tectonics t o Ore Deposits in t h e South Cordillera: Arizona Geological Socie ty Digest , v. XIV, p. 49-70.

Sillitoe, R. H., and Bonham, H. F., 1984, Volcanic landforms and o r e deposits: Economic Geology, v. 79, p. 1286-1298.

Ste iner , A., 1968, Clay minerals in hydrothermally a l t e r ed rocks a t Wairakei, New Zealand: Clay and Clay Minerals, v. 16, p. 193-213.

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Chapter 10 RELATIONSHIPS OF TRACE-ELEMENT PATTERNS TO GEOLOGY

IN HOT-SPRING-TYPE PRECIOUS-METAL DEPOSITS Byron R. Berger and Miles L. Silberman

INTRODUCTION

Those epi thermal precious-metal deposits where o re was precipi ta ted within 100-300 m of t h e earth 's surface such t h a t t h e d i rec t in teract ion of hydrothermal fluids with the surface is a major cause of ore-mineral precipitation in t h e upper pa r t of t he system make up t h e subclass known a s hot-s rin -t e deposits (Berger and Eimon, 1983; Berg* deposits were emplaced a s small veins, stockworks, and explosive breccias in association with non-marine volcanism, generally calc-alkaline in composition. Henley (198513, this volume) and Hayba et al. (1985, this volume) prefer t o no t s epa ra t e hot-spring deposits a s a s epa ra t e class or subtype of epi thermal deposits. However, we have chosen t o t r e a t hot-spring re la ted deposits separa te ly because of t h e importance of hydrothermal eruptions and accompanying brecciation to near-surface o re deposition and exploration recognition c r i t e r i a (Adams, 1985, this volume).

Act ive geothermal sys tems have long been thought t o be modern analogs of epi thermal sys tems (cf. White, 1955; Weissberg et al., 19791, but i t wasn't until t he r ecen t discovery of the McLaughlin gold deposit in California and t h e publication of d a t a on Round Mountain, Nevada (Berger and Tingley, 1980; Tingley and Berger, 1985) and Hasbrouck Mountain, Nevada (Silberman et al., 1979; Graney, 1984) t h a t t he re became a widespread recognition among explorationists of t h e geological and geochemical character is t ics and resource importance of fossil hot- spring systems. Subsequently, study in t h e Bodie, California mining dis t r ic t by P. Herrera and M. L. Silberman (Silberman and Berger, 1985, th is volume) has fur ther linked fossil hot-spring sys tems t o the deeper-emplaced bonanza-type epi thermal vein deposits by documenting a continuum of f ea tu res from the surface springs t o t h e deep veins. Recen t research a t Broadlands, New Zealand by K. Brown (Henley and Brown, 1985, th is volume) has shown t h a t t he geothermal fluids a r e nea r sa tura t ion with respect t o gold, and e lec t rum is ac t ively being precipi ta ted a t t h e wellhead. This implies t h a t alkaline-chloride thermal waters in general may b e sa tu ra t ed with gold, and t h a t t h e source of t h e gold is of less importance than the precipitation mechanisms. Although no economic concentrations of gold have yet been discovered a t Broadlands, r ecen t exploration in an ac t ive geothermal a r e a on Lihir Island, New Guinea has resulted in t h e discovery of a potentially economic gold deposit (R. Henley, D.S.I.R., personal communication, 1985). Thus the link between modern geothermal sys tems and epithermal-type vein deposits has been closed.

The t race-e lement chemistry of modern geothermal sys tems has been studied bv several Gorkers including Weissber (19691, ~ e i s s b e r ~ e t al. (19791, Ewers and Keays h977) , White (19811, and Henneberger (1983). These s tudies documented the high concentra t ions of arsenic, antimony, mercury, thallium, and tungsten in t h e upper par ts of geothermal sys tems in addition t o gold and silver. The t race-e lement distribution in relationship t o a l tera t ion in drill hole GS-5 a t Steamboat Springs, Nevada was summarized by Silberman and Berger (1985, this volume). Heretofore , similar geochemical studies of hot-spring-type precious-metal deposits have not been published. I t i s t h e purpose of this paper t o summarize t h e trace-element pat terns for gold, arsenic, antimony, mercury, and thallium in t w o well-studied hot-spring-type systems, Hasbrouck Mountain and Round Mountain, Nevada, and t o r e l a t e these pat terns t o t h e geology and hydrothermal a l tera t ion in e a c h system.

CONTROLS ON TRACE-ELEMENT PATTERNS

The system-wide trace-element pa t t e rns observed in o r e deposits represent t he summation of a multiplicity of processes t h a t a r e re la ted t o t h e t ime and space history of t he geothermal system, the variations in t h e fluid chemis t ry of t h e system, the chemis t ry of t h e host rocks, and t h e physical na tu re of t h e h e a t source and t h e hydrothermal system including fracturing, permeabili ty, and brecciation.

Silberman et al. (1979) documented (a) t h e length of t i m e t h a t hot-spring act iv i ty has been taking place at S teamboa t Springs, Nevada, and (b) t h e episodic na tu re of t h e geothermal ac t iv i ty (Table 10.1). They found t h a t hot-spring act iv i ty has been taking place in termit tent ly fo r more than 2 million years. Henneberger (1983) determined t h a t t he re have been four separa te , overlapping s tages of geothermal ac t iv i ty a t Ohakuri, New Zealand (Fig. 10.1). Henneberger showed t h a t t he location of t h e hydrothermal ac t iv i ty migrated with t ime. Sufficient t i m e elapsed between the s tages t h a t erosion took place between s t age 1 and s t age 2. Using ac t ive thermal sys tems both studies i l lus t ra te t h e t i m e and space complexities resulting in superimposed a l tera t ions and trace-element pat terns t h a t a r e probably present in a l l hydrothermal systems, including those t h a t a r e mineralized. Similar t iming and spat ia l variations in t h e physical aspects of hydrothermal sys tems was documented by Lloyd (1972) at Orakeikorako, New Zealand, where hot-spring act iv i ty

Table 10.1--Statigraphic and age relationships of geothermal activity and the emplacement of rhyolite intrusions and andesite flows in the SteamboaL Hot Springs geothermal field, Nevada (from Silberman et al., 1979)

Stratigraphic Unit Radiometric Age

- Recent alluvium; interbedded opaline sinter

- Lake Lahontan alluvium; interbedded opaline and cristobalite sinter

------- Erosional disconformity-------

- Pre-Lake Lahontan alluvium; interbedded cristobalite and chalcedonic sinter

- Interbedded, older chalcedonic sinter - Rhyolite pumice; intrusion of

Steamboat Hills Rhyolite

- Basaltic andesite - Rhyolite pumice; interbedded

chalcedonic sinter

- Alluvium; interbedded chalcedonic sinter

- Granodiorite (bedrock)

has resulted in a la rge number of s c a t t e r e d hydrothermal explosion cra ters , t h e foci of which changed with t ime. Ph rea t i c erupt ions occur over a r e a s where t h e h e a t flow is sufficiently high t o locally exceed t h e boiling point wi th dep th re la t ion (Elder, 1981). The implication a t Orakeikorako i s t ha t , a s a t Ohakuri, t h e ho t t e s t pa r t of t h e su r f ace hydrothermal ac t iv i ty has migra ted with t ime. Inact ive s in ter t e r r aces a t d i f ferent e levat ions and in d i f ferent a r eas f rom t h e presently ac t ive zone of fluid flow and s in ter development a t S t eamboa t Springs, Nevada, also document t h e migration of f luids through t i m e (White et al., 1964; Silberman e t al., 1979). Ti l ted s in ter and s in ter within some sedimentary deposits and beneath volcanic rocks a t S t eamboa t and younger sinter i l lus t ra te t h e long history of i n t e rmi t t en t ac t iv i ty with periods of erosion.

Henley (1985a,b; Henley and Ellis, 1983) reviewed geothermal fluid chemis t ry and t h e chemica l envi ronments of t ranspor t and deposit ion of precious metals. H e suggests t h a t t h e o r e and associa ted t r ace - e l emen t distr ibutions observed in epi thermal deposi t s resul t f rom t h e salinity and gas con ten t of t h e hydrothermal fluids. In turn, t h e t o t a l salinity of any given geothermal sys tem is dependent upon t h e host- rock compositions and t ec ton ic s e t t i ng (Henley, 1985b, th is volume). Basaltic- and silicic-hosted volcanic sys tems have lower salinit ies t han andesite-hosted systems. Gold-bearing sys tems have lower sa l in i t ies

and high sulfur contents whereas silver-base me ta l bearing sys t ems have re la t ive ly higher salinities. In t h e Taupo Volcanic Zone, New Zealand, Henley (1985a) has found a coincidence of high g a s flux ( C 0 2 + H2S) and gold, and he hypothesizes t h a t t h e gas i s derived f rom a deep source, thus requiring a s t ruc tu ra l s e t t i ng t h a t al lows upward flow f rom deep crus ta l (or even upper mantle?) regions.

The permeabili ty of t h e hos t rocks i s o n e of t h e mos t impor t an t physical a t t r i b u t e s a f f ec t ing t h e hydrothermal solutions, and i s t he re fo re a n impor t an t cont ro l on hydrothermal a l t e r a t ion a n d t race-e lement patterns. Unless hydrothermal sys t ems a r e smal l re la t ive t o t h e length a n d width of f rac tures , t h e permeabili ty will be f rac ture-contro l led a s was demonst ra ted by Elder (1981) in t h e Taupo, New Zealand volcanic zone. Porosity-controlled permeabili ty i s of increased impor t ance in smal l hydrothermal sys tems and locally in t h e larger systems. Faults, joints, and f r ac tu re s a r e a l l impor t an t in controll ing fluid flow, particularly when they const i tu te extensive, through-going s t ructura l systems. Porosity permeabi l i ty is a n in t r ins ic a spec t of e a c h individual host lithology. Sil tstones, shales, and welded tu f f s have low porosity permeabili ty whereas pumiceous tuf fs , tuf f breccias, and lapilli t u f f s a r e relatively permeable. Within any given s t ra t igraphic layer t h e porosity permeabi l i ty may vary result ing in s t ra tabound zones of re la t ive ly higher and

B. R. BERGER & M. L. SILBERMAN 235

Hydrothermal breccias a r e a n essent ia l a t t r i bu t e of t h e hot-spring subtype of epi thermal precious-metal deposits. They a r e commonly associa ted wi th high- grade o r e s (Berger and Eimon, 1983; Berger, 1985; Sillitoe, 1985; Tingley and Berger, 1985). Hedenquist and Henley (1985) classified hydrothermal eruptions a s e i ther shallow o r deep. Shallow erupt ions occur within a f e w m e t e r s of t h e su r f ace where s t eam flow is hampered by mineral deposit ion o r where t he re a r e changes in t h e near-surface hydrology. Deep hydrothermal eruptions occur below the s t eam condensate zone (Hedenquist and Henley, 1985) due t o hydraulic f rac tur ing (Grindley and Browne, 1976). The deeper breccias s e rve t o focus fluid flow, thus explaining t h e impor tant association of breccias with o r e in hot-spring-type deposits. Hedenquist and Henley (1985) hypothesize t h a t hydrothermal eruptions occur where t h e r e i s localized overpressuring in t h e upper 300 m of t h e geothermal system. The local sealing of fluid-flow paths by minera l deposition causes sl ight increases in fluid pressure with fluids

Figure 10.1. The location of distinct stages of hydrothermal activity at Ohakuri, New Zea- land, based upon stratigraphic and alteration studies by Henneberger (1983). Early activity was characterized by zones of quartz-adularia, silicification, and zeoli- tization. Intermediate activity is recognized by the addition of kaolinite alteration probably resulting from a change in the water table. Late-stage activity is characterizedby silicification and a diminishing volume of hydrothermal activity.

lower permeability. The physical na tu re of t hese zones may a lso change with t ime. The deposit ion of minerals (e.g., quar tz , ca lc i te , zeoli tes, etc.) m a y sea l porous zones, thereby reducing permeabili ty. Subsequent hydrothermal f rac tur ing or t ec ton ic ac t iv i ty may induce fracture-controlled permeabili ty, reopening "locally sealed" pa r t s of t h e system.

When t h e sys tem consists of two-phase flow, such a s a vapor-liquid mixture, t h e flow of o n e phase a f f e c t s t h e flow r a t e of t h e o the r phase. Dif ferent ia l flow r a t e s c a n result in varying vapor-liquid ra t ios within t h e system; these ra t ios a r e impor tant cont ro ls on t h e parti t ioning of t r a c e e l emen t s be tween t h e liquid and t h e vapor. Therefore, t race-e lement pa t t e rns may b e dependent upon t h e vapor-liquid r a t i o in t h e fluid a s well a s t h e gas composit ion of t h e vapor phase. As t h e deeper-heated wa te r s near t h e surface , processes occur which modify t h e fluid flow (Elder, 1981). Permeabi l i ty var ies la tera l ly due t o su r f ace f rac tures , and topography becomes a n impor t an t influence. A subsurface boiling wa te r t ab l e may resul t in s t eaming ground where e x t r e m e leaching a n d acid- su l f a t e a l t e r a t ion assemblages may be ad j acen t t o a l te ra t ion produced by alkaline-chloride waters. Pa t chy a r e a s of hot ground may be surrounded by cold ground due t o t h e downward percolation of cold ground wa te r (Elder, 1981). All of t hese su r f ace processes resul t in complex trace-element distr ibution pat terns , part icularly when combined with temporal migra t ion of t h e t he rma l ac t iv i ty a s discussed previously.

predomina;tly flowing t o o ther unblocked flow channels. Vapor accumula t e s in t h e sea led f r ac tu re s and gas exsolution causes increased pressure on t h e sea l f rom t h e gas cap. Hydraulic f r ac tu r ing will s e t off a hydrothermal eruption--the f r ac tu r ing e i the r due t o t h e gas c a p exceeding l i t hos t a t i c pressure o r t ec ton ic ac t iv i ty reopening t h e sea led pathways.

TRACE-ELEMENT PATTERNS IN STUDIED DEPOSITS

Hasbrouck Mountain, Nevada

Hasbrouck Mountain is located in t h e Divide mining d is t r ic t about t h r e e miles south of t h e town of Tonopah in west-central Nevada. Cold was f i rs t discovered on Cold Hill just e a s t of Hasbrouck Mountain in 1902, and smal l gold mines in t h e d is t r ic t produced in t e rmi t t en t ly until high-grade silver o r e was discovered in 1917 below t h e mined levels on Cold Hill (Bonham and Garside, 1979). The g r e a t e s t period of production in t h e d is t r ic t was between 1920 and 1929. The Kernic vein on Hasbrouck Mountain was unusual in t h e ea r ly history of t h e d is t r ic t in t h a t i t was mined largely f o r i t s si lver con ten t ins tead of gold. The exploited mineralization on Hasbrouck Mountain was r e s t r i c t ed t o narrow veins t h a t o f t e n consisted of si l ica-cemented breccias, and the re i s no record of any significant production. To ta l gold-silver production in t h e d is t r ic t was about $3.5 million (Bonham and Carside, 1979). In t h e 1970ts, Hasbrouck Mountain was explored fo r i t s low-grade, large-tonnage gold potential , and a n o r e body of about 3 million tons averaging about 0.05 ounces Au/ton was discovered (A. Wallace, Cordex Exploration, wr i t t en communication, 1982). Subsequent geologic s tudies on Hasbrouck Mountain (Silberman et al., 1979; Bonham and Carside, 1979; Graney, 1984) showed t h e o r e sys tem t o be re la ted t o hot-spring ac t iv i ty a s evidenced by t h e presence of chalcedonic s in ter and extens ive , multiple episodes of hydrothermal breccia t ion , including a n erupt ion breccia in terbedded with Sieber t Format ion

Epiclastic volcanic sandstone; local plant fossils; epiclastic volcanic conglomerate localized hypothermal eruption breccias

Chalcedonic sinter

Pumiceous ash-flow tuff Rotary drill hole and hole number

Water-lain tuff, laminated siltstone, ash-flow tuff, ' sandstone and conglomerate f Normal fault

Fraction tuff: Tonopah Summit member A * * 8 a n n A t Cross-section

Figure 10.2. The generalized geology of Hasbrouck Mountain, Nevada, based upon the work of Graney (1984). A-A' is the line of the northwest-southeast cross section shown in Figure 10.3, with the dots showing the number and location of rotary drill holes. The top of Hasbrouck Elountain is approximately located southeast of RDH-5 along the cross-section line near the northeast- trending faults.

Volcan~c cgl-ss

Chalcedon>c slnter

_______----- _____------ ______-------

Tuff, lamnated slltstone. Conglomerate and sandstone

Figure 10.3. A northeast-looking, generalized geologic cross section of Hasbrouck Moun- tain, Nevada, also showing the locations of rotary drill holes for which alteration and trace element data are given in Figures 10.4 a-e. Refer to Table 10.2 for more detailed stratigraphic information.

sandstones (R. Henley, D.S.I.R., personal communication, 1984).

The a rea l geology of t h e Divide dis t r ic t was described by Bonham and Garside (1979) and t h e geology of Hasbrouck Mountain by Graney (1984). A simplified version of t h e geologic map of Hasbrouck Mountain is shown in Figure 10.2, and a cross section is shown in Figure 10.3. The oldest s t ra t igraphic units a r e t h e Tonopah Summit and King Tonopah members of t h e Fract ion Tuff, a qua r t z l a t i t e ash-flow tuff (Table 10.2). Silberman et al. (1979) repor t potassium-argon ages of 18.2 t o 19.9 m.y. on bioti te and alkali-feldspar from the tuff. The Fract ion Tuff i s overlain by t h e middle Miocene Sieber t Formation, a composite unit of in tercala ted sedimentary, volcaniclastic, and volcanic rocks which yield K-Ar ages of 16.2 t o 15.5 (Silberman e t al., 1979). Rhyol i te domes and re la ted dikes of approximately t h e s a m e age intrude t h e older units. Bonham and Garside (1979) repor t two episodes of dome emplacement. One is t h e Oddie rhyolite which is consanguinous with t h e mineralization in the district; t h e second being t h e post-mineralization Brougher domes. A number of northerly trending normal faul ts t ransect Hasbrouck Mountain. The stratigraphy a s well a s t h e s t ruc tu re const i tu tes an important control on t h e t race-e lement distribution patterns. The l o w e r ~ n o s t Siebert unit consists of fluvial sandstones and conglomerates with some ash-flow tuff beds. Water-lain tuffs, ash-flow tuff, epic las t ic volcanic si l tstone, volcanic conglomerate, and sandstone form a sequence t h a t underlies chalcedonic sinter. Volcanic sandstone, si l tstone, conglomerate, and ash-flow tuff a r e interbedded with and overlie t h e sinter and make up the youngest Sieber t Formation exposed on Hasbrouck Mountain. The ash-falls and ash-flows a r e commonly poorly welded and some contain abundant pumice lapilli and therefore a r e not as easily f rac turable a s t h e welded tuffs and thinly laminated water-lain tuff. Also, the poorly welded mater ia l has a relatively higher porosity permeabili ty than t h e welded and thinly laminated rocks. Hydrothermal ac t iv i ty was taking place locally a t Hasbrouck Mountain during t h e development of t he volcanic field

hosting t h e activity. K-Ar ages of hydrothermal adularia and ser ic i te from Hasbrouck a r e in t h e range of 16.3 t o 15.3 m.y. (Silberman et al., 1979). Local volcanic units, including the rhyolite plugs and other flow units, a r e in t h e range of 16.8 t o 14.8 m.y.

The deta i led and ca re fu l study by Graney (1984) outlined t w o discre te foci of gold mineralization based upon t h e distribution of gold and the a l tera t ion patterns. These t w o foci appear a s "mushroom- shaped" o re zones with concentr ic a l tera t ion envelopes. We in t e rp re t t hese o r e zones a s t h e s i t e s of t h e primary conduits along which the hydrothermal fluids were flowing. Figures 10.4a through e show t h e hydrothermal a l tera t ion in a f ence of drill holes t h a t t r ansec t s t h e t w o mineralized conduits. RDH-5 (Fig. 10.4a) and RDH-2 (Fig. 10.4b) pene t r a t e the cen t r a l portion of e a c h conduit. Quartz-adularia a l tera t ion gives way t o a n assemblage of quartz-adularia-albite a t depth. The quartz-adularia zones in t h e t w o foci coalesce in t h e upper par t of t he system rendering the appearance on t h e surface of only one relatively large a l tera t ion zone. RDH-21 (Fig. 1 0 . 4 ~ ) was collared in quartz-adularia-altered rock and went in to quartz- adularia-illite a l tera t ion beneath it. The pa t t e rns in drill holes 2, 5, and 21 suggest t h a t in t h e main fluid conduits two-feldspar a l tera t ion marks the axial root of t h e ore-bearing quartz-adularia zone, and t h a t both t h e a lb i t ic and adularia zones a r e flanked by i l l i t ic alteration. RDH-9 (Fig. 10.4d) and RDH-20 (Fig. 10.4e) a r e peripheral t o t h e main conduit a r eas and show a marked increase in i l l i t ic a l tera t ion and t h e appearance of montmoril lonite a s an a l tera t ion product. The d a t a presented by Graney (1984) indicate t h a t i l l i t ic a l tera t ion is also present a s a major a l tera t ion mineral beneath t h e adularia- and albite- bearing zones.

Exposures in underground workings of t h e quartz- adularia zone show i t t o be extensively brecciated. The breccias may b e smal l or very extensive, and appear t o have happened repeatedly in single locations. There a r e two major classes of brecciation-- uncemented and cemented. The uncemented breccias consist of rock f ragments of varying sizes in a finely comminuted matr ix of rock flour and rubble. These breccias a r e in terpre ted t o have resulted f rom sudden eruptions of locally over-pressured vapor-enriched fluids with t h e brecciation resulting from t h e rapid expansion of t h e gases upon breaching of t h e localized sealing and consequent boiling. The continuity of finely laminated beds across this type of breccia indicates t h a t t he re has been l i t t l e or no ver t ica l t ranspor t of rock f ragments during t h e explosive eruptions. The cemented breccias a r e similar in t ex tu re and s t ruc tu re t o t h e uncemented var ie ty excep t t h a t t h e rock f ragments a r e supported by a matr ix of si l ica and disseminated pyrite. The o r e occurs a s quartz-adularia veins and veinlets within and around t h e breccias a s well a s within t h e silica- cemen ted breccias.

The t race-e lement pat terns observed a t Hasbrouck Mountain a r e complex, and in a l l likelihood a r e due t o the in teract ion of several physical and chemical processes taking place episodically within t h e fluid conduits. As a consequence of this colnplex genesis, in deta i l t he distributions of individual

0 Mercury, ppm

Manganese, ppm

Gold, ppm

Mercury, ppm ~m

VArsenic and +antimony, ppm loo 200 300 a

Quartz-adularia

m~uar tz -adu la r ia

.Thallium, ppm

RDH-2vArsenic and *antimony, ppm 0 100 200 300

0 2 4 6 8 Thallium, pprn

*Antimony and thallium, ppm

Manganese, pbin

Mercury, ppm

RDH-2 1

.Thallium and *Antimony, ppm e

VArsenic, ppm

Quartz-illite

Quartz-illite- montmorillonite

Quartz - adularia - albite - illite

Quartz - illite

Quartz - adularia

Quartz - adularia - illite

Figure 10.4.a-e. Hydrothermal alteration and data for selected trace elements in rotary drill holes along cross section A-A' (refer Figs. 10.2 and 10.3).

a). Data from RDH-5 which penetrated the central part of the eastem-most focus of hydrothermal activity. The heavy black line from 50-70-foot depth shows the location of chalcedonic sinter (refer to Fig. 10.3) in the hole. The heavy black line from 220-240-foot depth marks an interval of intense brecciation.

b). Data from RDH-2 which penetrated the central part of the western focus of hydrothermal activity.

c). Data from RDH-20 which penetrated the most distal alteration zone away from the main focus of hydrothermal activity.

d). Data from RDH-9 which penetrated lateral alteration zones adjacent to the main focus of hydrothermal activity.

e). Data from RDH-21 which penetrated a small lateral wedge of quartz-adularia altered rock carrying ore-grade gold and bottomed in a peripheral alteration zone.

Table 10.2--Stratigraphic relationships at Hasbrouck Mountain, Nevada (after Graney, 1984; Bonham and Garside, 1979; ages from Silberman et al., 1979)

- Tertiary opaline sinter (some plant fragments)

- Unaltered Tertiary air-fall ash -------------- Main Mineralizing Event---------------- (15.3-16.3 m.y.1

- Miocene Divide Andesite; quartz latite dike (approximate age 14.8-16.9 m.y.1

- Miocene Siebert Formation (15.5-16.2 m.y.1

'Volcanic conglomerate; some interbedded sandstone and ash-flow tuff

'Volcanic sandstone and siltstone; some plant fragments

'Chalcedonic sinter

'Volcanic conglomerate with some sandstone; local hydrothermal fall-back breccias

'Pumiceous ash-flow tuff

'Water-lain tuff; interbedded volcanic siltstone and sandstone

'Crystal-lithic ash-flow tuff

'Fluvial sandstone and conglomerate

- Miocene Fraction Tuff

'King Tonopah Member; vitric-lithic rhyolite ash-flow tuff (approximate age 18.2-19.9 my.)

'Tonopah Summit Member; vitric quartz latite to rhyolite ash-flow tuff

elements and t h e relationships between the various e l emen t s a r e themselves complex and vary depending on t h e scale a t which the system is evaluated. Figures 10.4a-e show t h e ver t ica l distributions of se lec ted t r a c e e l emen t s in a f ence of drill holes r e fe r r ed t o above. Assuming t h a t t h e foci of t h e hydrothermal system a r e outlined by the distribution of quartz- adular ia a l tera t ion, t h e concentrations in drill cut t ings of gold, silver, arsenic, antimony, mercury, and thallium a r e relatively more enriched in t h e potassically a l t e red centra l pa r t s of t he system (Table 10.31, while zinc, lead, manganese, and boron a r e enriched in a l l a l tera t ion zones peripheral t o the quartz-adularia a l tera t ion zones. When t h e concentra t ions of e lements in t h e quartz-adularia- i l l i te zones (peripheral t o t h e quartz-adularia) a r e compared t o the more distal quartz-illite- montmoril lonite zones, thallium, mercury, and z inc a r e more enriched in the outer quartz-illite-

montmoril lonite zones and gold and arsenic a r e relatively more abundant in the quartz-adularia-illite zones. Therefore, a t t h e scale of t h e whole system, t h e r e appears t o be an inner zone of arsenic, antimony, thallium and mercury re la ted t o t h e precious-metal o re and an outer zone of mercury, boron, and manganese t h a t bounds t h e o r e zones. The results a r e reminiscent of t h e e lementa l zoning surrounding the Au mineralized zone a t Paramount (Silberman and Berger, 1985, this volume). Although o r e occurs only in association with quartz-adularia a l tera t ion, gold is above t h e analytical detect ion level throughout t h e a l t e red area.

The t race-e lement distributions in t h e drill holes shown in Figures 10.4a-e a r e complex when examined in detail. RDH-5 (Fig. 10.4a) was collared stratigraphically higher than t h e chalcedonic s in ter (shown a s heavy black l ine from 60-70-foot depth on Figure 10.4a) and penetra ted t h e en t i r e breadth of t h e

Table 10.3--The concentration of selected trace elements associated with the quartz- adularia zones at Hasbrouck MounLain, Nevada, vis-a-vis concentrations associated with all other alteration zones

Quartz-adularia zone (n = 193 samples) All other alteration zones (n = 205 samples)

Element Minimum Maximum Element Minimum Maximum ( P P ~ ) ( P P ~ ) ( P P ~ ) ( P P ~ )

ore-bearing quartz-adularia a l tera t ion zone. The t o p of t h e higher-grade gold zone occurs about 80 f e e t beneath t h e sinter. Mercury, antimony, and arsenic mimic t h e gold distribution pattern. Manganese also appears t o be higher in t h e gold zone in th is hole. The heavy black line between 220-240-foot depth in RDH-5 (Fig. 10.4a) marks an interval of in tense brecciation. Manganese shows a marked decrease within t h e breccia, thereby effect ively forming a halo around the breccia zone. Manganese oxides a r e commonly observed adjacent t o breccias both on the surface and in underground workings. As a general rule gold, arsenic, and antimony show a weak negat ive corre la t ion with depth in both RDH-5 and -2 (Fig. 10.4b), t h e centra l conduits of t h e hydrothermal system. Mercury and thallium show no particular t rend with depth in these holes. Peripheral t o t h e main pa r t of t he hydrothermal system (RDH-21, -9, and -20, Figs. 10.4~-el, t h e gold, arsenic, and antimony continue t o decrease with depth a s does thallium t o a sl ight degree, but t h e mercury tends t o show a n increase with depth.

Round Mountain, Nevada

Round Mountain is located in t h e Round Mountain mining dis t r ic t about 50 miles nor th of t he town of Tonopah, Nevada, on t h e western slope of the Toquima Range. Cold was initially discovered a t Round Mountain in the 18908s, but production did not t ake place until 1906, and significant lode mining continued until 1942. Through 1969, about 350,000 ounces of gold and 360,000 ounces of silver were produced (Tingley and Berger, 1985). In t h e 1960's renewed exploration outlined substantial tonnages of low-grade gold-silver mineralization containing 9.4 million ounces gold and 16.9 million ounces silver, and open-pit mining has been ongoing since 1977. Based upon geologic and geochemical information gathered by Tingley and Berger (1985) a t Round Mountain, t h e o r e deposit was in terpre ted by them t o have been formed in t h e upper portion of a hot-spring system. Their evidence includes a breccia pipe with included

blocks of sandstone, conglomerate, and s in ter 100-200 f e e t down in t h e pipe beneath thei r original s t ra t igraphic level, a hydrothermal-eruption (?) breccia interbedded with epic las t ic rocks and syn- sedimentary quartz-sulfide beds within t h e sediments.

The regional geology around Round Mountain was mapped by Shawe (1981), and t h e deta i led deposit geology (Fig. 10.5) by Tingley and Berger (1985). A schemat i c s t ra t igraphic column is shown in Figure 10.5. The oldest rocks in t h e d is t r ic t a r e complexly folded Cambrian t o Ordovician schist , quar tz i te , argil l i te, and l imestone intruded by a Cre taceous muscovite-biotite granite. In t h e Fairview mine (Fig. 10.5), t h e grani te is pervasively a l t e red with milky qua r t z veins. L a t e Oligocene t o ear ly Miocene rhyolite ash-flow tu f f s a r e the main host rocks t o ore. The tuffs yield K-Ar ages of 26.1 t o 24.7 m.y. on biot i te and sanidine. The lowest unit is a welded ash- flow tuff megabreccia of undetermined thickness. Blocks of Paleozoic sediments and Cre taceous grani te occur in t h e breccia and vary in s ize from pebbles t o blocks several hundred f e e t across. A complex, thick quartz-sanidine rhyolite ash-flow unit overlies the megabreccia consisting of a lower 400- t o 500-foot- thick non-welded base grading upward in to a n 800- foot-thick welded tuff which is then transit ional upward into a 75- t o 100-foot-thick non-welded top. The ash-flows a r e capped by a sequence of epic las t ic rocks including water-lain tuff, sandstone, conglomerate, and bedded silica ( in terpre ted by Tingley and Berger (1985) t o be subaqueous chalcedonic sinter).

Cold mineralization is controlled primarily by northwest-trending f r ac tu res (Fig. 10.6) and occurs a s low-angle and ver t ica l vein and shee t zones, breccia- filling and stockworks in t h e welded tuff , and a s microveinlets and disseminations in the lower non- welded tuff immediately above the megabreccia. All of these mineralization types a r e genetically r e l a t ed and their location and variation in form mainly r e f l ec t variations in physical properties of t he host rocks. The o r e minerals a r e e lec t rum, f r e e gold, and auriferous pyrite. K-Ar age determinations on adularia and

A

2 E a, C

a 6

Metasedimentary rocks

Figure 10.5. A generalized geologic map and schematic stratigraphic column for Round Mountain, Nevada (from Tingley and Berger, 1985). The sample locations for data used to construct Figures 10.7-10.10 are shown as black dots. The open-pit outline shows the mine as it was in late 1979. The area of bedrock outcrop is shown in a stippled pattern.

ser ic i te within veins and adjacent t o breccias range from 25.1 t o 25.5 m.y. (Tingley and Berger, 1985).

All of t h e veins conta in qua r t z and adularia and have irregular, but persistent, a l tera t ion envelopes of initially sericit ic-argil l ic and then propylitic mineral assemblages. The gold occurs a s f r e e gold and e lec t rum on t h e qua r t z and adularia crys ta ls and within sulfides in t h e veins. The phyllic-argillic a l tera t ion ref lec ts t h e presence of plagioclase totally replaced by white mica (a mixture of ser ic i te and illite), b io t i te complete ly a l t e red t o white mica and chlorite, and pumice replaced by mixtures of white mica and kaolinite. Sanidine phenocrysts a r e embayed. The propylit ic a l tera t ion consists of plagioclase partially a l t e red t o white mica, bioti te variably a l t e red t o chlorite-white mica, and t h e rock matr ix a l t e red t o a clay-white mica mixture.

The breccia zones and stockworks a r e common, but only one a r e a in t h e Sunnyside mine beneath Round Mountain has been mined t o any extent . The breccias formed f rom t h e r epea ted fracturing, quar tz veining and sil ica flooding, and rebrecciation of t h e host tuff. Vuggy qua r t z o f t en cemen t s t h e breccia fragments. Gold occurs a s f r e e gold and e lec t rum and within sulfides in t h e vugs and a s inclusions in sulfides in t h e silica matr ix of t h e breccia. A zone of intense phyllic- argillic a l tera t ion surrounds t h e silicified breccias.

A disseminated, blanket-like deposit occurs in t h e lower non-welded portion of the ore-bearing ash- flow tuff , and t h e a l tera t ion in this zone consists of quartz, white mica, and possibly adularia. The mineralization occurs a s microveinlets of quartz, sulfides filling pumice cavities, and disseminated sulfides throughout t h e tuff. Quartz-chlorite a l tera t ion generally underlies the ore-bearing rock.

The distribution of mineralization a t Round Mountain is complex and cannot be simplified in to the distinct and sepa ra t e cen te r s of hydrothermal ac t iv i ty in the s a m e manner t h a t Craney (1984) was able t o do a t Hasbrouck Mountain. Nevertheless, t he re is a general zone beneath Round Mountain and Stebbins Hill defined by pervasive silicification centered about Round Mountain and even more extensive phyllic- argillic a l tera t ions t h a t accompany vein and breccia ore, and a second zone on t h e southwestern s ide of Round Mountain where veins t ransect ing propylitically a l t e red rock predominate and t h e phyllic-argillic a l tera t ion is confined t o narrow selvages along the veins (Fig. 10.6). The a r e a of pervasive a l tera t ion (hachured pa t t e rn on Figure 10.6) was in terpre ted by Tingley and Berger (1985) t o be a linear chain of closely spaced fluid channels and the a l tera t ion envelopes around t h e individual channels have coalesced t o c r e a t e t h e apparent single a l tera t ion zone. The channels themselves a r e marked by the presence of hydrothermal breccias, each breccia representing multiple even t s of cementat ion with quartz, adularia, and sulfides and subsequent rupturing. O r e i s found both within sulfides and vugs in t h e breccias and in sulfides along low-angle veins t h a t in general dip towards t h e breccia cen te r s and appear t o be concentr ic around the breccias. Ore- bearing high-angle veins a r e found t o both crosscut t h e breccias and be t runca ted by the breccias. Tingley and Berger in t e rp re t t h e low-angle veins t o have been

B . R. BERGER & M. L. SILBERMAN

F i g u r e 10.6. G e n e r a l i z e d s t r u c t u r e a n d a l t e r a t i o n map o f Round Mountain, Nevada (mcdif i e d f rom T i n g l e y a n d Berge r , 1985). The p e r v a s i v e areas o f s i l i c i f i c a - t i o n a n d p h y l l i c a l l y al tered r o c k o c c u r w h e r e the re a re discrete f o c i o f h y d r o t h e r m a l brecciat i o n i n the subsur face . Q u a r t z - a d u l a r i a v e i n s o c c u r a l o n g a l l o f the n o r t h w e s t - t r e n d i n g f r a c t u r e s and some o f the n o r t h e a s t - t r e n d i n g f rac- t u r e s .

derived from the multiple injection of vein const i tuents along low-angle cooling joints in t h e welded tuff during t h e hydrothermal explosions t h a t c rea t ed t h e breccia centers. The sequential injection of mater ia l has dilated the low-angle joints t o produce t h e thick veins; t h e vein thicknesses decrease away from the inferred cen te r s of fluid flow and brecciation. The second zone of mineralization probably represents severa l separa ted fluid channels and a long, l inear hydrothermal vent a rea known a s the Automat ic structure.

As a consequence of the complex mineralization history a t Round Mountain, t he geochemical pat terns (Figs. 10.7-10.10) have an even more complicated expression and genesis than those at Hasbrouck Mountain. In general, a r eas t h a t a r e enriched in gold and silver (Fig. 10.7) a r e also anomalous in the pathfinder e lements arsenic and antimony (Fig. 10.8, thallium and mercury (Fig. 10.91, and molybdenum and tungsten (Fig. 10.10). The gold deposits a r e not character is t ica l ly enriched in copper, zinc, and lead, although locally these base meta ls occur in anomalous amounts. Zinc and copper a r e most commonly a t highest concentra t ion along with silver in a reas of manganese enrichment. The mineralized a r e a character ized by pervasive a l tera t ion is marked by a n enr ichment of arsenic, antimony, and thallium and a

Alunlte p r e s e n t a s velns

S l l l ~ l f l c a t l o n e l ther pervasive or a s par t la1 r e p l a c e m e n t o f r o c k phenoc;y919 a n d l o r groundmans

P e r v a l l v e alteration 01 tuft b lo t l te to chlorlte andlor ser ic i te Secondary c l a y s

PrODyl8t~zatlOn b lot l te a l t e r e d to chlortte a n d l o r rer lc l te

depletion of calcium and manganese. Tungsten, si lver, and molybdenum a r e also enriched in this a r e a with t h e tungsten r e s t r i c t ed t o t h a t par t of t h e a r e a with t h e most in tense silicification. Gold is present in amounts exceeding 0.05 par ts per million throughout th is a r e a of pervasive a l tera t ion (Figs. 10.6 and 10.7), but also overlaps in i t s ent i re ty t h e adjacent zone on t h e southwest side of Round Mountain and is thus an indicator e l emen t of t he whole breadth of t h e highly mineralized area . There is a small-scale geochemical zoning t h a t i s discernible around e a c h of the cen te r s of brecciation in t h e pervasively a l t e red zone. Manganese oxides occur in abundance along t h e low- angle veins ou t away f rom t h e breccias. The manganese is seen in some instances t o completely ring breccia pipes and pods. As mentioned, t h e manganese is enriched in si lver and the base metals. Gold along t h e low-angle veins and within sulfides in t h e breccias averages about 650 fine, whereas crystall ine gold in vugs within t h e breccias is g rea t e r than 900 fine.

The second a l tera t ion zone on t h e southwest s ide of Round Mountain contains anomalous arsenic, antimony, and thallium within t h e veins, but t he re is no pervasive sa tura t ion of large concentra t ions of these e l emen t s ou t in to t h e propylitcally a l t e red wallrock. Very high-grade concentrations of f r e e gold

Figure 10.7. The distribution of gold and silver in rock samples from the surface at Round Moun- tain, Nevada (from Tingley and Berger, 1985).

Figure 10.8. The distribution of arsenic and antimony in rock samples from the surface at Round Mountain, Nevada (from Tingley and Berger, 1985).

were s toped a t t w o locations, t h e nor thwest (Sphinx P i t ) and southeas t (No. 2 Glory Hole) ends of t h e Au toma t i c s t ruc ture . Near t hese high-grade zones t h e Au toma t i c s t ruc tu re is intensely b recc i a t ed a n d argil l ized and t h e f r ac tu re s a r e commonly c o a t e d wi th fluorite. Away f rom t h e Au toma t i c s t ruc ture , p,er se, breccias a r e not extens ive suggesting t h a t s l ~ g h t l y d i f ferent physical processes occurred during mineralization in this zone a s compared t o t h e pervasively a l tered zone.

Arsenic, antimony, mercury, and thall ium appea r t o have t h e highest concentra t ions in s u r f a c e samples in t h e s t ruc tura l ly high pa r t s of t h e hydrotherinal

sys tem (Figs. 10.8 and 10.9). Antimony is highest in a r e a s of more intense argi l l ic a l t e r a t ion such a s on t h e e a s t side of Stebbins Hill. In order t o document t h e changes in m e t a l concentra t ions wi th depth, i t was necessary t o analyze a se r i e s of ro t a ry drill holes across both of t h e genera l ized a l t e r e d areas. The appa ren t dec rease in su r f ace samples in t h e more volati le e l emen t s with dep th i s also demons t r a t ed f rom t h e analyses of drill cutt ings. The lowest whole-rock concentra t ions of arsenic and ant imony occur in t h e blanket-like deposit in t h e lower non-welded tuf f , al though se lec ted samples of qua r t z microveinlets in th is deposit t ype a r e st i l l highly anomalous in arsenic.

Figure 10.9. The distribution of mercury and thallium in rock samples from the surface at Round Mountain, Nevada (from Tingley and Berger, 1985).

Figure 10.10. The distribution of tungsten and molyWenum in rock samples from the surface at Round Mountain, Nevada (from Tingley and Berger, 1985).

Although thallium decreases with depth in t h e lower non-welded tuf f , t h e dec rease in concentra t ion is less d rama t i c t han with e i ther a r sen ic o r antimony.

DISCUSSION

The major goals in studying t h e t race-e lement pa t t e rns a t Hasbrouck Mountain and Round Mountain a r e t o provide insights in to t h e origin of ore-bearing hot-spring sys tems and t o provide d a t a and in terpre ta t ions useful t o explorationists. Based upon t h e deta i led study of t hese t w o deposits, i t was found

t h a t t hey have a number of f e a t u r e s in common. Both of t h e deposits studied concen t r a t ed similar su i tes of elements--gold, si lver, arsenic, antimony, mercury, thall ium, molybdenum, and tungsten. Both deposits have a paucity of base m e t a l s associa ted with t h e ore. Additionally, both a r e r e l a t ed t o non-marine rhyolit ic volcanic ac t iv i ty with evidence of consanguineous high-level in t rus ive ac t iv i ty , and a t both deposits t h e t iming of minera l iza t ion overlaps t h e period of ac t ive volcanism. T h e hydrothermal fluids appear t o have been focused i n t o speci f ic channels with repeated episodes of localized self-sealing through minera l deposit ion followed by brecciation and

subsequent o r e deposition. I t is no t known if t ec ton ic ac t iv i ty , hydrothermal overpressuring due t o a gas c a p a s sugges ted by Hedenquist and Henley (19851, or both, caused t h e breccia t ion o r ruptur ing of t h e seal. The occurrence of b recc i a ve ins and dikes ostensibly unre la ted t o through-going f r ac tu re s sugges ts t h a t overpressuring plays s o m e role in hydrothermal erupt ions although a c t i v e tec tonism i s known t o accompany volcanic a n d in t rus ive activity.

F o r t h e mos t pa r t , ore-grade concentra t ions of t h e precious m e t a l s a r e associa ted with t h e prec ip i ta t ion of t h e gangue minera ls qua r t z and adularia. This assemblage is deposited in t h e main hydrothermal flues. T h e pathfinder e l emen t s arsenic, antimony, thall ium, and mercury a lso have t h e highest concentra t ions within t h e main fluid channels and a r e t hus e f f ec t ive indica tors of t h e foci of hydrothermal activity. A t Hasbrouck Mountain and Round Mountain, mercury is anomalous bo th within t h e co re potassic a l t e r a t ion zone and in t h e outer periphery of t h e hydrothermal system. Manganese shows a similar pa t t e rn a s mercury by forming a peripheral ring around t h e main a reas of fluid flow. Silberman and Berger (1985, th is volume) documented Hg and Mn halos around individual ve ins o r mineralized zones in similar pa r t s of minera l ized sys tems a t Paramount and Bodie. Thus, t he se e l emen t s give evidence of being more generally useful f o r delineating a r e a s of mineralization. However, t h e Bodie d a t a and di f ferences in t h e distribution of t hese e l emen t s in t h e four hydrothermal sys t ems described show t h a t considerable variation occurs.

Within t h e co re potassic zone, t h e ver t ica l t race- e l emen t pa t t e rns appea r t o be dependent upon t h e re la t ive permeabi l i t ies of t h e rocks. More porous portions of sedimentary beds in t h e Sieber t Format ion a t Hasbrouck Mountain show increased concentra t ions of a l l t h e pathfinder e l emen t s (Figs. 10.3 and 10.4). Areas of breccia t ion of both Hasbrouck Mountain a n d Round Mountain a lso a r e enr iched in t hese s a m e t r a c e elements.

La t e ra l f low m a y become particularly impor tant when higher e levat ions flank t h e upper pa r t s of t h e sys tems and cold su r f ace wa te r s fo rce t h e fluid t o flow outward (Hanaoka, 1980) or when t h e fluids flow la tera l ly beneath a n impermeable layer. In s i tua t ions where l a t e r a l f low h a s t aken place, a zone of argil l ic a l t e r a t ion underlies t h e main fluid channel, thus giving t h e appea rance of a l a t e r a l ra ther than a n underlying a l t e r a t ion assemblage. La te ra l flow leads t o hot- spring ven t a r e a s t h a t a r e up t o severa l k i lometers away f rom t h e main h e a t source a s a t Waiotapu in New Zealand ( ~ e d e n q u i s t and Henley, 19851, and R. Henley (D.S.I.R. , personal communication, 1983) specula tes t h a t t h e Fairview mine a t Round Mountain may be a l a t e r a l f low ven t area . The a l t e r a t ion a t t h e Fai rv iew is cons is tent with th is in terpre ta t ion in t h a t in tense quartz-white mica a l tera t ion underlies t h e quartz- adular ia ore-bearing rock. A similar in terpre ta t ion may be m a d e of t h e Hasbrouck Mountain mineralization in t h a t i l l i t ic a l t e r a t ion underlies t h e potassic zone with t h e "feeder" s t ruc tu re s dipping t o t h e e a s t towards Gold Hill where t h e r e has been a focus of intrusive activity.

The compilation of geochemical d a t a f rom

Hasbrouck Mountain and Round Mountain shows t h a t t h e physical and a l t e r a t ion f ea tu re s of t h e t w o deposi t s a r e of considerable impor t ance t o t h e pa t t e rns observed. Deposit-specific s t ruc tu ra l f e a t u r e s a n d host-rock permeabi l i t ies appea r t o be very impor t an t cont ro ls on the focusing of hydrothermal fluids i n to individual f lues in each deposit. Therefore , geochemical and a l t e r a t ion zoning will vary la tera l ly and ver t ica l ly be tween di f ferent deposits and n o theore t ica l t race-e lement pa t t e rns o r speci f ic concentra t ions of se lec ted pathfinder e l emen t s should b e expected . However, geochemical zoning can b e determined f o r any given sys tem o n a n empir ica l basis and may then b e used a s a guide t o mineralization in t h a t d is t r ic t . Elements t h a t s e rve a s d i r ec t pathfinders t o t h e precious me ta l s in one deposit m a y be less useful in ano the r deposit , al though most hot-spring-type minera l deposits probably conta in t h e s a m e genera l su i t e of elements--gold, silver, a rsenic , antimony, thall ium, mercury , tungsten, and molybdenum. Occasionally, e l emen t s such a s selenium and tellurium may be impor t an t pathfinder e l emen t s (e.g., DeLamar , Idaho), al though they were no t found t o be of impor tance in t h e two deposits discussed in th is paper.

REFERENCES

Adams, S. S., 1985, Using geological informat ion t o develop exploration s t r a t eg i e s f o r epi thermal deposits; & Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemis t ry of Epi thermal Systems: Socie ty of Economic Geologists, Reviews in Economic Geology, v. 2.

Berger, B. R., 1985, Hot-spring type gold deposits; Tooker, E. W. (ed.), Geologic Cha rac t e r i s t i c s of Sediment and Volcanic-Hosted Disseminated Gold Deposits--Search f o r an Occurrence Model: U.S. Geological Survey, Bulletin 1646.

Berger, B. R., and Eimon, P. L., 1983, Conceptual models of epi thermal precious-metals deposits; in Shanks, W. C., I11 (ed.), Cameron Volume on - Unconventional Mineral Deposits: Socie ty of Mining Engineers, p. 191-205.

Berger, B. R., and Tingley, J. V., 1980, Geology and geochemis t ry of t h e Round Mountain gold deposit , Nye County, Nevada (abs.): Prec ious Metals Symposium, Nevada Bureau of Mines and Geology, Reno, NV, p. 18c.

Bonham, H. F., Jr., and Garside, L. J., 1979, Geology of t h e Tonopah, Lone Mountain, Klondike, and nor thern Mud Lake quadrangles, Nevada: Nevada Bureau of Mines and Geology Bulletin 93, 142 p.

Elder, J. W., 1981, Geothermal Systems: Academic Press, London, 508 p.

Ewers, G. R., and Keays, R. R., 1977, Volatile- and precious-metal zoning in t h e Broadlands geothermal f ield, New Zealand: Economic Geology, v. 72, p. 1337-1 354.

Graney, J. R., 1984, Controls of a l t e r a t ion and precious-metal mineralization in a fossil hydrothermal sys tem, Hasbrouck Mountain, Nevada (abs.): Geological Society of America , Abs t r ac t s with Programs, v. 16, no. 6, p. 523.

Grindley, G. W., and Browne, P. R. L., 1976, St ructura l and hydrological fac tors controll ing t h e permeabili ty of some hot water geothermal fields: United Nations Symposium Development and Use Geothermal Resources, 2nd, San Francisco, Proceedings, p. 377-386.

Hanaoka, N., 1980, Numerica l model exper iment of hydrothermal system-topographic ef fec ts : Bulletin of t h e Geological Survey of Japan, v. 31, p. 321-332.

Hayba, D. O., Bethke, P. M., Heald, P., and Foley, N.K., 1985, Geologic, mineralogic, and geochemical cha rac t e r i s t i c s of volcanic-hosted epi thermal precious-metal deposits; & Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemis t ry of Epi thermal Systems: Society of Economic Geologists, Reviews in Economic Geology, v. 2.

Hedenquist, J. W., and Henley, R. W., 1985, Hydrothermal eruptions in t h e Waiotapu geothermal sys tem, New Zealand: Their origin, associa ted breccias , and relation t o precious- me ta l mineralization: Economic Geology, v. 80, p. 1640-1668.

Henley, R. W., 1985a, O r e t ranspor t and deposition in epi thermal environments: Proceedings Symposium on Stable Isotopes and Fluid Processes in Mineralization, University of Brisbane, Geological Society Australia Special Publication, p. 1-43.

Henley, R. W., 1985b, The geothermal f ramework of epi thermal deposits; & Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemistry of Epi thermal Systems, Society of Economic Geologists, Reviews in Economic Geology, v. 2.

Henley, R. W., and Brown, K. L., 1985, A pract ica l guide t o t h e thermodynamics of geothermal fluids and hydrothermal o r e deposits; Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemis t ry of Epi thermal Systems, Society of Economic Geologists, Reviews in Economic Geology, v. 2.

Henley, R. W., and Ellis, A. J., 1983, Geothermal systems, anc i en t and modern: Ea r th Science Reviews, v. 19, p. 1-50.

Henneberger, R. C., 1983, Petrology and evolution of t h e Ohakuri hydrothermal sys tem, Taupo volcanic zone, New Zealand: Unpublished M.S. thesis, University of Auckland, 141 p.

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Silberman, M. L., and Berger, 1985, Relationship of t race-e lement pa t t e rns t o a l tera t ion and morphology in epi thermal precious-metal deposits; & Berger, B. R., and Bethke, P. M. (eds.), Geology and Geochemis t ry of Epi thermal Systems, Socie ty of Economic Geologists, Reviews in Economic Geology, v. 2.

Silberman, M. L., Bonham, H. F., Jr., Garside, L. J., and Ashley, R. R., 1979, Tirning of hydrothermal alteration-mineralization and igneous ac t iv i ty in t h e Tonopah mining d i s t r i c t and vicinity, Nye and Esmeralda counties, Nevada; & Ridge, J . D. (ed.), Proceedings of t h e Fi f th Quadrennial Symposium, International Association on the Genesis of O r e Deposits: Nevada Bureau of Mines and Geology, Repor t 33, p. 119-126.

Silberman, M. L., White, D. E., Keith, T. E. C., and Doctor, R. D., 1979, Duration of hydrothermal ac t iv i ty of S t eamboa t Springs, Nevada f rom ages of spatially associa ted volcanic rocks: U.S. Geological Survey, Professional Paper 458-D, 13 p.

Sillitoe, R. H., 1985, Ore-re la ted breccias in volcanoplutonic arcs: Economic Geology, v. 80, p. 1467-1514.


Weissberg, B. G., 1969, Gold-silver ore-grade prec ip i ta tes f rom New Zealand the rma l waters: Economic Geology, v. 64, p. 95-108.

Weissberg, B. G., Browne, P. R. C., and Seward, T. M., 1979, O r e m e t a l s in a c t i v e geothermal systems; in Barnes, H. L. (ed.), Geochemis t ry of - Hydrothermal O r e Deposits, Second Edition: John Wiley and Sons, New York, p. 739-780.

White, D. E., 1955, Thermal springs and epi thermal o r e deposits; & Bateman, A. M. (ed.): Economic Geology, 50th Anniversary Volume, p. 99-154.

White, D. E., 1981, Act ive geothermal sys t ems and hydrothermal o r e deposits: Economic Geology, 75th anniversary volume, p. 392-423.

White, D. E., Thompson G. A., and Sandberg, C. A,, 1964, Rock s t ruc tu re and geologic history of S t eamboa t Springs t he rma l area , Washoe County, Nevada: U.S. Geological Survey, Professional Pape r 458-B, 6 3 p.

Chapter 11 BOILING, COOLING, AND OXIDATION IN EPITHERMAL SYSTEMS:

A NUMERICAL MODELING APPROACH Mark H. Reed and Nicolas F. Spycher

INTRODUCTION

Some ac t ive geothermal sys tems a r e current ly depositing gold, silver, and base meta ls , and most "epithermal" o re deposits formed in once-active geothermal systems (e.g., White, 1981; Henley, 1985, th is volume). Boiling of hot ( 1 0 0 ~ - 3 0 0 ~ C ) ground water in such sys tems is a process of fundamental significance because i t f ixes t empera tu re gradients (e.g., White et al., 1971; Muffler e t al., 1971; Henley and Ellis, 1983) and causes precipitatiorr of sulfide, carbonate, and s i l ica te minerals (e.g., Buchanan, 1981; Berger and Eimon, 1983). The gas phase, including H20 , C 0 2 , and H2S, when condensed and oxidized near t h e surface , produces acid waters t h a t genera te argil l ic a l tera t ion of rocks and which may trigger deposition of precious metals. The geologic and hydrologic framework of a boiling geothermal system is depicted in Figure 11.1, based in pa r t on White e t al. (19711, Henley and Ellis (19831, Berger and Eimon (1983), and Steven and Eaton (1975). Figure 11.2 corresponds t o Figure 11.1, showing in flow-diagram form t h e chemical components and processes in the hydrothermal system. These include boiling (A, Figs. 11.1 and 11.2)) condensation of the boiled gas in rock (B), oxidation of the gas by the a tmosphere (C), condensation followed by oxidation of t h e gas in cool, f rac tured ground (D), mixing of ac id ground waters with t h e boiled liquid (E), and mixing of cold ground water with the boiled liquid (F). All of t hese processes shape t h e chemistry of geothermal sys tems and severa l of them a r e responsible for ore format ion in epi thermal systems. We present he re some results of detailed calculations of heterogeneous equilibria (Reed, 1982) t h a t apply t o t h e boiling, cooling, fluid- fluid mixing, condensation, and oxidation depic ted in Figures 11.1 and 11.2. Our focus here is on epi thermal o re formation, so most details of t h e calculation approach in general and many deta i ls of reaction chemistry a r e reserved for publication elsewhere (Reed and Spycher, in preparation).

The results presented here a r e an extension and ref inement of ear l ier work (Reed and Spycher, 1983) on boiling of a dilute, gas-rich water based on t h a t f rom Broadlands Well No. 2 (Mahon and Finlayson, 1972). In a re la ted study using homogeneous equilibrium calculations on numerous geothermal waters, Reed and Spycher (1984) showed how the pH and mineral solubilities in the aqueous phase a r e influenced by boiling and dilution. More recently, Drummond and Ohmoto (1985) used homogeneous and partial heterogeneous equilibrium calculations to evaluate t h e e f f e c t s of boiling on mineral precipitation in epi thermal systems. Drummond and Ohmoto (1985)

provide a deta i led discussion of several important f ea tu res of boiling sys tems t h a t we omi t here, including: open- vs. closed-system boiling; kinetics of redox reactions; and some e f f e c t s of major e lement solution chemis t ry on t h e s i ze of pH changes caused by boiling. Our calculations a r e fo r a water resembling t h a t from Broadlands Well No. 2 (Mahon and Finlayson, 19721, reconst ructed (Reed and Spycher, 1984) f rom analyses of t he liquid and gas a t 25OC (Table 11.1). Trace-metal contents fo r th is water a r e modified from Weissberg e t al. (1979). Af t e r completing t h e study presented here, we learned t h a t t h e long-accepted values for gold concentra t ion in t h e Broadlands water e r r in being too low (Henley and Brown, 1985, this volume) thus, with respect t o gold, our water does not ma tch t h e ac tua l Broadlands water a s closely a s we intended. The results are , nevertheless, valid in thei r own right, applying t o a Broadlands-like water with a low initial gold concentration. The initial water composition given in Table 11.1 is in equilibrium numerically with chlorite, galena, muscovite, and sphalerite a t 27S°C, resulting in slightly decreased molali t ies of severa l components re la t ive t o the Broadlands water. A Broadlands-like water was se lec ted fo r init ial study fo r several reasons: (a) a complete analysis i s available, including t r a c e meta ls and gases, (b) i t i s from a well-studied hydrothermal system, providing abundant field d a t a for comparison with our results (Browne and Ellis, 1970; Mahon and Finlayson, 1972; Ewers and Keays, 1977; Grant, 1977; Weissberg e t al., 19791, and (c) Broadlands waters have large C 0 2 , H2S, and CH4 gas fractions; thus, the chemical e f f e c t s of degassing should be distinct.

For the purpose of exploring the geochemical processes i l lustrated in Figure 11.2, we s t a r t ed by boiling 1.015 kg of the init ial solution (1 kg of H 2 0 plus 15 grams of dissolved species). All successive

operations on t h e liquid plus gas system (condensation, mixing, oxidation, etc., Fig. 11.2) discussed and il lustrated below were e f f e c t e d without re-normalizing t o a unit of gas or unit of liquid. Thus, a l l extensive quantit ies presented below a r e "per 1.015 kilogram of init ial unboiled fluid."

BOILING

When a homogenous aqueous phase rises in a hydrothermal sys tem (Fig. 11.1, A), i t experiences a decrease in pressure resulting in boiling. Boiling induces a t empera tu re dec rease and a pH increase which causes minerals t o precipitate. Thus, t he initially homogenous aqueous phase separa tes in to several phases including gas, liquid, and minerals. The

Table 11.1--Boiling with fractionation

Temperature (deg.C.) 101 151 202 230 242 250 262 278 278

Pressure (bars) 1.02 4.69 Gas wt. % 36 28

PH 7.80 7.31

Component species (total molalit~)

H+

c1-

F-

so4-- so4--( ion) cog-- HS-

Si02

Al+++

~ a + +

M ~ + +

Fe++

K+

~ a +

~ n + +

cu+

~ b + +

Ag+

A"+

Hg++

Gases (mole % )

Minerals (moles/deg.C.)

Clinochlore

Daphnite

Galena .492E-11 .308E-10

Sphalerite .303E-13 .247E-12

Quartz

Microcline .165E-09 .316E-08

Calcite

Chalcopyrite

Bornite .334E-14 .314E-12

Chalcocite .622E-10

Acanthite .244E-10

M. H. REED & N. F. SPYCHER 25 1

Figure 11.1. Schematic diagram of a boiling hydrothermal system based in part on White et al. ( 1 9 7 1 ) , Henley and Ellis ( 1 9 8 3 ) , Berger and Eimon ( 1 9 8 3 ) , and Steven and Eaton (1975) . Ascending hot waters begin to boil (indicated by circles) at the depth labeled H, if pressure is hydrostatic or at some shallower depth between H and L if pressure exceeds hydrostatic but is less than lithostatic. Boiling and various mixtures of boiled hot water and gases with near- surface ground water and atmospheric oxygen produce hot- water compositions and ore- forming environments labeled by letters A through F. These reaction features are identified by letter and described in Figure 11.2 and in the text.

Ac~d Sulfate Waters

f

- H

(and metal leaching)

Magmat~c Heat Source at Depth

gas, liquid, and some of the minerals a r e themselves solutions. A t equilibrium in any such mixture of specific bulk composition a t pressure (PI and t empera tu re (T), overall heterogeneous equiliibrium can be calcula ted (Reed, 1982) yielding: (a) t h e activit ies, molalities, and distribution of aqueous species; (b) t h e masses, identit ies, and compositions of minerals; and (c) t he mass and composition of the gas phase. In t h e results presented here, such calculations a r e carr ied o u t for stepwise decreases in P and T along a predetermined, nearly isoenthalpic boiling curve (explained below). By this means we s imulate the chemical e f f e c t s of boiling during t h e ascent of a geothermal water initially a t 2 7 8 ' ~ and 67.8 bars t o t h e surface where T = 100°C and P = 1 bar.

The calculation of the results presented he re depa r t s f rom the approach of Reed (1982) only in t r ea t ing t h e gas phase a s ideal ra ther than non-ideal. Eight components, H20, C 0 2 , H2S, CH4, HCI, HF, H2, and S2, a r e used t o represent t he a s phase. Where oxidized gases a r e important far in oxidation calculations below), 0 2 is included also. Thus, a l l chemically significant species excep t NH3 (Giggenbach, 1980) a r e included. The inclusion of economically in teres t ing but chemically insignificant gases composed of Hg, As, and Sb is t h e object of fu tu re work.

Solid solutions of chlorite and white mica were t r ea t ed using ideal multi-site mixing (Kerrick and Darken, 1975). All Henry's Law constants for gas- liquid parti t ioning a r e implicit ly accomodated in t h e equilibrium constants for t h e appropriate mass ac t ion equations (equation (1 2) of Reed, 1982). Dependence of Henry's Law constants on salinity (e.g., Ellis and Colding, 1963; Drummond, 1981) is implicit ly accounted fo r in t h e ac t iv i ty coefficients for neutra l aqueous species (e.g., Helgeson, 1969).

Isoenthalpic boiling is one common and useful assumption for t h e behavior of geothermal fluids

(Elder, 1965; Grant e t al., 1984; Henley e t al., 1984, p. 9, 143; Drummond and Ohmoto, 1985). We determined a nearly isoenthalpic equilibrium a t many t empera tu res by traversing pressure from high t o low a t e a c h temperature . By calculating t h e enthalpy of t h e gas-liquid mixture along e a c h isothermal pressure t raverse , we determine a pressure a t which t h e enthalpy equals t h a t of t h e pure liquid a t t h e t empera tu re and pressure of incipient gas separation (278OC and 67.8 bars for our water). This fixes a ser ies of P-T points along an isoenthalpic boiling curve and i t provides a wealth of information concerning super- and sub-isoenthalpic cooling t ra jec tor ies . The l a t t e r information is presented separa te ly in a following section.

A be t t e r way t o determine an isoenthalpic boiling curve would be t o incorporate an enthalpy equation in t h e system of simultaneous equations, having specified an initial enthalpy. Then, t h e isoenthalpic P-T curve could be determined by successsively decrement ing P (or T ) and solving for T (or P) t h a t sa t i s f ies t h e enthalpy equation. We a r e current ly developing this capability.

Having fixed a P-T boiling curve, t he heterogeneous equilibrium calculations a r e carr ied out by stepping down t h e P-T curve in increments of 5' t o 16OC. For boiling along such a curve, t w o cases a r e t rea ted: (a) boiling with mineral f rac t ionat ion and (b) boiling a t constant bulk composition. A third possibility, boiling with gas f rac t ionat ion (Rayleigh distillation, e.g., Drummond and Ohmoto, 1985) with or without mineral fractionation, is also of in teres t , but w e have not ye t ca lcula ted it. Boiling in a sys tem t h a t is closed with respect t o gas separa t ion allows continuous chemical communication between t h e liquid and t h e gas separa t ing f rom i t a s they ascend in a f r ac tu re system. This appears t o b e a reasonable f i rs t approximation of rea l systems.

Boiling with mineral f rac t ionat ion is execu ted by

BOILING RESULTS

Boil: isoenthalpk P-decrease

Parent Uquid - 28WC

Figure 11.2. Flow diagram showing some of the boiling, mixing, and reaction processes that a deep geothermal water experiences. The processes depicted here, labeled A through F, correspond to those with the same labels in Figure 11.1 and described in the text. The notations along the arrows indicate the type of calculation presented here for the corresponding processes. Boiling (A) is presented in Figures 11.3, 11.4, and 11.7; oxidation of condensate and gas (C) is in Figure 11.10; mixing of gases with ground water (D) is in Figure 11.11; mixing of acid-sulfate water with boiled water (E) is in Figures 11.12 and 11.13; mixing of ground water with boiled water is in Figure 11.14.

subtrac t ing f rom t h e bulk composit ion before each P-T dec remen t a n amoun t of t h e various chemical components corresponding t o those which were fixed in minerals on t h e just-finished P-T step. This corresponds in na tu re t o a n ascending boiling fluid "leaving behind" t h e minera ls prec ip i ta ted f rom it. They a r e thus not avai lable t o back-react a t lower P and T. In t h e constant bulk composit ion case, a l l mineral prec ip i ta tes a r e kep t in chemical con tac t with t h e solution throughout boiling. They a r e available t o back-react and t o const ra in t h e solution chemistry. Another ca se discussed below is a "cooling only" calculation, fo r which t h e pressure is held high enough throughout incrementa l cooling t o prohibit format ion of a gas phase. This corresponds t o cooling ent i re ly by conduction and provides a n ins t ruct ive comparison t o t h e boiling calculations.

The thermochemical d a t a fo r t hese calculations a r e compiled f rom diverse sources, including Helgeson (1969); Helgeson and Kirkham (1974a, b; 1976); Walther and Helgeson (1977); Helgeson et al. (1978); Reed and Spycher (1984); Wolery (1979); Barnes (1979); and Schwar tzenbach and Widmer (1966). All of t h e d a t a fo r complexes of me ta l s with sulfides and bisulfides (e.g., Pb, Zn, Cd, Hg, Au) a r e f rom t h e pr imary references given by Barnes (1979). Fo r mos t of t hese da t a , we co r r ec t ed t h e constants t o a s tandard s t a t e of infinite dilution using a modified Debye-Huckle equat ion (Helgeson e t al., 19811, and we ext rapola ted measured values t o cover t h e t empera tu re range f rom 300°C t o 100°C.

Calcula ted resul ts of boiling with mineral f rac t ionat ion a r e given in Figures 11.3a-i and Table 11.1. Resul ts fo r closed-system boiling a r e depic ted in Figures 11.4a-d. Cooling-only results a r e shown in Figures 11.5a-d. These Figures a r e bes t read from r ight t o le f t , corresponding t o t h e a scen t of fluids in t h e geothermal system.

I t is convenient in t h e following discussion t o use t h e t e rm, "early" t o r e f e r t o chemical even t s toward t h e high-temperature end of t hese plots. Also, r e f e rence t o "rapidity" of mineral precipitation or molali ty changes r e f e r s t o a r a t e of change with decreas ing t e m p e r a t u r e a n d corresponds t o t h e slope of curves in t h e graphs. In a l l cases, 1.015 kilograms of fluid were used a t t h e s t a r t ; thus, results expressed in grams o r moles r e f e r t o quant i t ies of gases or minerals per 1.015 kilogram of init ial liquid. Since minerals were f r ac t iona t ed in t h e calculation shown in Figure 11.3, t h e minera l F igures 11.3b and 11.3d refer t o quant i t ies prec ip i ta ted per deg ree of t empera tu re change (per 1.015 kilogram of init ial fluid); i.e., they represent a f in i te-d i f ference approximation of t h e der ivat ive of minera l abundance with respect t o t empera tu re , whereas t h e minera l f igures in 11.4 and 11.5 r e f e r t o t h e absolute mass of minerals present. Qua r t z disappears a t 200°C in Figure 11.3b because i t was arbi t rar i ly suppressed a t t h a t point t o s imula te i t s fa i lure t o mainta in equilibrium with t h e aqueous phase a t low t empera tu re s (Fournier, 1985, th is volume; Reed, 1985).

DISCUSSION O F BOILING AND COOLING

The profound e f f e c t of boiling on t h e hydrothermal chemis t ry c a n bes t be apprecia ted by compar ing t h e boiling calculations t o a simple cooling calculation (Fig. 11.5) on t h e s a m e water composition. In t h e cooling-only case, pH decreases (Figs. 11.5b and 11.6) a s weak acids (H2C03, HCL, HSOG) dissociate with decreas ing tempera ture . In t h e boiling reactions, pH increases, a s discussed below. The influence of this pH e f f e c t re la t ive t o t empera tu re change, among others , is explored below by compar ing t h e boiling calculation resul ts t o t h e cooling-only results.

Figure 11.3a shows t h e composition of t h e gas phase continuously in equilibrium with t h e aqueous phase throughout boiling. (Note t h a t t h e corresponding plot fo r t h e closed-system boiling calculation closely resembles Figure 11.3a, and is omitted.) Scrutiny of curves f o r t h e predominant "dry" gases, C 0 2 , CH4, HZS, and H , shows t h a t t hese gases a r e rapidly f r ac t iona t ed Trom t h e a ueous phase in t h e f i rs t 5OC 8 t o 2 5 ' ~ of boiling (278 -250'~). Below 250°c, t h e abundance of H 2 0 gas s teadi ly increases while t h e dry gases change l i t t le. This ear ly degassing e f f e c t divides t h e boiling i n t o t w o s t a g e s which makes a dist inct imprint on t h e solid phase assemblage. The ear ly s t age of rapid degassing of C 0 2 causes a n abrupt increase in pH (decrease in a ~ + , F ~ g s . 1 1 . 3 ~ and 11.4b) a s t he

following reaction is displaced by C 0 2 escape f rom t h e aqueous phase

(Note, here and below, an arrow is used t o indicate a reaction forced in the indicated direction by the boiling or cooling process occurring over a P-T interval. An equal sign is used for equilibria a t specific P-T points.) Degassing of aqueous H2S

could also contr ibute t o pH increase, but since t h e quantity of H2S is qu i t e smal l (Fig. 11.3a) re la t ive t o C 0 2 , i t s influence is minor. Thus, i t is primarily reaction 1 t h a t results in a change in pH from 6.4 a t 27S°C t o 7.0 a t 245OC t o 7.8 a t IOOOC (Fig. 11.3~) .

In contras t t o t h e major gases, t h e absolute amounts of HCI and HF in t h e a s phase reach maxima 5 in t h e range of 225OC t o 250 C (Fig. 11.3a). Their early rise is a ref lec t ion of removal of molecular HF and HCI present in t h e aqueous phase a t high temperatures , e.g.

With decreasing temperature , t he d ie lec t r ic constant of water increases (Helgeson and Kirkham, 1974a), stabilizing ionic species, leading t o dissociation of aqueous HCI which displaces t h e above reaction back t o t h e left .

Sulfide and Carbonate Mineral Precipitation

Figures 11.3b and 11.4a show t h a t t h e highest r a t e s of mineral precipitation prevail over t h e short t empera tu re interval between 2 7 8 ' ~ and 2 4 5 ' ~ where the t empera tu re dec rease combines with t h e pH increase t o fo rce the following example sulfide and carbonate reactions

zn2+ + HS- Z ~ S + H+ ( 3 ) ( a q ) ( a q ) ( s p h a l e r i t e ) ( a q )

The greater significance of pH increases t o these reactions re la t ive t o t empera tu re decrease is apparent from the absence of ca lc i te , sphalerite, chalcopyrite, and galena a t high t empera tu re in t h e cooling calculation (Fig. 1 l.5a), re la t ive to t h e boiling calculations (Figs. 11.3b and I l.4a). The more acid conditions in t h e cooling-only case (Fig. 11.6) displace reactions 3 and 4 t o t h e left , inhibiting mineral precipitation. For sulfides, this occurs despite the increase in metal-ion molali t ies (Fig. 1 1 . 5 ~ ) due t o dissociation of thei r chloride and sulfide complexes (Fig. 11.5d).

Over t h e s a m e shor t t empera tu re interval (278'- 2 4 5 ' ~ ) where most mineral mass precipitates, t he loss

of t h e reduced gases, H2 and CH4,.causes oxidation of aqueous sulfide t o sul fa te , i n c r e a s ~ n g sul fa te molali ty a s shown in Figure 11.3i. Degassing also causes to t a l aqueous carbonate and sulfide (represented a s CO; and HS-, Fig. 1 1 . 3 ~ ) t o decrease while t h e molali ty of individual ion CO? increases (Fig. 11.3i), because t h e increase in pH displaces t h e following react ion t o the right

Individual ion HS- decreases during boiling (Fig. 11.3f) due t o loss of sulfide t o t h e gas phase, but t h e dec rease is subdued re la t ive t o the change in to t a l molali ty of sulfide (Fig. 1 1 . 3 ~ ) because t h e increase in pH results in dissociation of aqueous H2S

Transfer of moles of H2S to t h e gas phase (Fig. 11.3a) ent i re ly dominates over sulfide mineral precipitation (all i n the range of moles o r less, Fig. 11.4a) in removal of sulfide f rom t h e aqueous phase. This is par t ly a consequence of low initial base- me ta l contents in t h e solution, consistent with i t s low salinity (and, therefore , low C1- content ; s e e fur ther discussion below).

Despite separa t ion of sulfide in to the gas phase, t h e high-sulfide concentration in this water makes sulfide and bisulfide complexes dominant (Figs. 11.3h, 11.4d, and 11.5d) fo r a l l o re meta ls excep t Zn, for which a bisulfide is secondary t o a chloride, and Pb, for which a carbonate is dominant over a shor t interval. All of t h e complexes become less s table with decreasing temperature , which tends t o re lease t h e me ta l s t o form sulfide minerals. Breakup of bisulfide complexes due t o HZS escape in boiling fur ther promotes me ta l p rec ip~ ta t ion , despite the moderating e f f e c t of H2S dissociation, reaction (6).

Figure 11.3d, which shows the re la t ive r a t e of precipitation of sulfide minerals with decreasing temperature , indicates t h a t sphalerite and galena precipi ta te f rom this water before chalcopyrite. The ear ly precipitation of sphalerite results from t h e combined e f f e c t s of increasing pH (reaction 3) and dissociation of ZnC1' (Fig. 11.3h) with decreasing temperature . Figures 11.3b and 11.3d show t h a t t h e r e i s a gap in galena precipitation between 2 4 5 ' ~ and 200 '~ . This i s a consequence of a shift in t h e dominant P b complex ligand from HS- t o C 0 5 and back again t o HS- (Fig. 11.3h). The carbonate complex t akes over temporarily a s mCo5 increases (Fig. 11.3f) owing t o increasing pH (reaction 5).

The l a t e precipitation of chalcopyrite (af ter sphaler i te and some galena) shown in Figure 11.3d is a consequence of i t s very stable bisulfide complex (Fig. 11.3h). Henley e t al. (1984) points ou t t h a t Broadlands waters precipi ta te more chalcopyrite t han sphaler i te or galena l a t e in the boiling process.

Because t h e water is sulfide rich, t he re i s suff ic ient sulfide t o stabil ize the high-sulfur copper minerals, bornite and chalcopyrite, over chalcopyrite with decreasing t empera tu re (Fig. 11.3d), a s discussed

254 CHAPTER 11

Boi l wi th Frac . Gases C _ _ _ _ _ _ - - - ---2Z--, @

- , - t , , , , , , , , , I , , , , : , , , ,

__ - - - - - k Boil with Frac . 92 I

Figure 11.3. Calculated isoenthalpic boiling of 1.015 kg of a Broadlands-like water with fractionation of minerals: a). Moles of species in the gas phase, continuously in equilibrium with minerals and the aqueous phase; b). Rate of precipitation of minerals in moles per degree of temperature change; c). Total molality of a port ion of the aqueous component species (not individual ion molality). Activity of hydrogen ion is also shown, for which the scale refers to log activity; d). Paragenetic-style diagram showing relative rate of sulfide mineralprecipita- tion on a weight basis; e). Total molality of o e metal component species; Fe3+ -refers to total aqueous Fe.

1 ' ' ' ' ' ' ' ' ' 1 ' ' , ~ , ' ' ' ' '

~ a + CI- ~?a- - ' r= - - - - -+ - -+- -z -= ---- -.-.-----, ---- -Fs"J2 c

----/ LK+

C _ _ C _ _ _ _ _ C _ _ _ _ _ - - - - - - - - - - - - - 9s ca2+

Al3+ //----c-----

Boi l wi th Frat. Component Species

1 . 1 . . . . . . . 1 , . , # . . , , , 1 . , , , , . , , ,

l ' ' ' ' r ' ' ' ' l ' r ' l l ' ~ ~ ' I ' ' ' ~ ' ' ~ ' '

Boi l wi th Frac . acan Minerals

bn

A * - - gn -gn @

1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 I I , I I , , , , , ,

Component Species 1

Figure 11.3 (cont'd.) f). and g). Individual species molalities of component species; h). Indi- vidual species molalities of the principal ore-metal complexes; i). Individual species molalities of carbonate, sulfate, and sulfide species. Mineral abbreviations used here and in the following Figures are as follows: acanthite, acan; bornite, bn; calcite, calc; chalcocite, cc; chalcopyrite, cp; chlorite, chl; cinnabar, cinn; clinochlore, clchl; covellite, cv; daphnite (Fe- chlorite), daph; graphite, graph; hematite, hem: kaolinite, kaol; microcline, K-sp; muscovite, musc; pyrite, py; quartz, qz; sphalerite, sl.

l ' ' ' ' r r r ' r l ' ' l ' ' ' ' ' ~ I ' ~ ' ~ ' ~ ' l '

~ a + CI- S%---- 0

K+

---------- -------- es- - - - - - - - - - - F-

co 2- ---------- -.---L------ >i2+ --- --,QH+ -- . - - - - -,,--, ~rr-::z--I:ZZZ~~~ ---'02*

Boil with Frac. Component Species

1 " " " " ' 1 " " " " ' ~ " " " " ' ~

: Boil with Frac. -

-

-

- @:

1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 (

1 " " " " ' l " " " ' " I " " " " ' .

:Boi l with Frac. Metal c o m p l e x e ~ ~ + : - -

- 12 100 150 200 250 300

Temperature CDe8.C.)

1 " " " " ' I " " " " ' I ' " " " " .

-Boi l - with Frac. Acidic ----.--*-------.-------------- - H C O ~ a= - I -'-;.rrc=---c----- F@-

HS --- - % - - - ----__ ---.- co32-

soi2- - - -----_._-_ - ----- - _-----------------_---- - ---*--- =-c----i~+ -

.-- 4- -- - a ~ + 1. '-- - - -

~ ~ 0 4 - - 1 . . . . . , # , , I . . . . . . . . . l . , . . . . . . .

Boi l w i th0u.t Frac. Minerals -.--- --------

qz -------_____ ---.

!

- 0 -6 M- -.-----T 4 0 I*, , , , , , , , -7---- , , , , , , , , , ;;;, , , , - , ST?==$ ;; , , , @I 4 -

OD -8 Boi l without Frac. 0 A Component Species -10

100 150 200 250 300

Temperature CDeg.C.)

Figure 11.4. Calculated isoenthalpic boiling of a Broadlands- like water without f ractiona- tion of minerals. The gas phase composition is essentially the same as in Figure 11.3a. a). Moles of minerals present in equilibrium with the gas and aqueous phases during boiling. See caption to Figure 11.3 for mineral abbreviations. b). Total molality of a portion of the component species. Activi- ty of hydrogen ion is also shown, for which the ordinate refers to log activity. c). Total molality of o e-metal mmpcnent species. Felt refers to total aqueous Fe. d). Indi- vidual species molalities of principal complexes of ore metals.

Figure 11.5. Cooling of a Broad- lands-like water without boiling and without mineral fractionation. (a) Moles of minerals present in equilibrium with the aqueous phase. See caption for Figure 11.3 for mineral abbreviations. (b) Total molality of a portion of the component species. Activity of hydrogen ion is also shown, for which the ordinate refers to log activity. (c) Total molality of ore-metal component species. (d) Individual species molalities of principal complexes of ore metals.

l l l l l ' ' l " l l l l l l l l l ' l l l ' ' l r l ' '

Cool ing Only* Minera ls ---------- --- - - W h - - - - _ - @

l l ' l l ' ' ' l l l l l l l ' ' l ' ' l ' l ' ' T r ' ' l

~ a + CI-

-----7-=-= 7-7-T -- Si02 K+ - -

ca2+

2+ - - - - - - - - - -w- - - - - - - ------.-__-___ 2- -\- s o ~ - - . - -. - -. - - ~h~:---------------,.-.~;------ - - - --- - - _ _ . ~ a ~ ~ - .

Component Spec i es '--.Qz- 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1

1 ' ~ ~ ~ " " ' l " ~ ~ ~ ~ ~ " 1 " ~ ~ ~ " ~ '

Cooling Only Component Species 0; - F.~+ -

-

- - +,=" -

/ -- -*+- - - - -

/ . , , , l . , , - r rT ; . . . . . . . , . I . , , , , , , . , :

150 200 250 300

Temperature CDeg.C.1

L " " " " ' ~ ' " " " " ~ " " " " ' ~ " " " ~ "

- CuIHS12- ---- __--------- --_______ -- -

. AgHS - - - -

Cooling Only - Metal Complexes

1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 3 . 1 1 -

Temperature CDeg.C.)

F i g u r e 11.6. Comparison of pH i n c a l c u l a t i o n s of boi l ing (lower curve, b t h boi l ing calcula t ions) and c o o l i n g o n l y (upper curve). See t e x t fo r discussion of differences.

by Reed (1985). The high-sulfide concentration also holds Ag' in the stable A HS complex (Fig. 11.3h) 8 throughout cooling from 278 to 115O, where acanthite precipitates (Figs. 11.3b and 11.3d). In contrast to silver, gold and mercury never precipitate because their sulfide and bisulfide complexes (Fig. 11.3h) are quite stable over the entire temperature and pH range of boiling. Gold and mercury behavior are more thoroughly discussed below.

Precipitation of Silicates

In both boiling calculations, solid solutions of chlorite and muscovite precipitate as well as quartz and K-spar (Figs. 11.3b and 11.4a). Solid solution compositions (calculated using ideal and ideal multi- site mixing) are not plotted separately, but compositions can be read from the figures by comparing the moles of solution end members. The calculated chlorite is iron rich, reflecting the high Fe/Mg ratio of the starting solution (Table 11.11, and the rapid dumping of Fe from solution during cooling as Fe(OHI4- dissociates.

In contrast to the sulfides and carbonates, the very presence of silicates is not a consequence of the pH increase upon boiling. The effects of temperature decrease control silicate precipitation, as is obvious for quartz but perhaps not so obvious for the others. Muscovite and K-spar precipitate because A1 is liberated from AI(OHl4- as its stability decreases with decreasing temperature (Reed and Spycher, 1984). This displaces the following example reaction to the right

Chlorite precipitates in response to this A1 supply combined with Fe supplied by dissociating Fe(0H);. The validity of these conclusions is indicated by the nearly quantitative removal of aqueous A1 and Fe in both the boiling reactions and in the cooling-only reaction (Fig. Il.5a), where pH decreases (Figs. 11.5b and 11.6). The high pH of the boiling waters fixes feldspar rather than mica in accordance with the following reaction

This reaction also accounts for the replacement of mica by feldspar with decreasing temperature (and increasing pH) shown in Figures 11.3b and 11.4a.

Boiling Without Fractionation and Cooling Only

As discussed above, most of the chemical processes in the closed-system boiling calculation are the same as in the boiling with mineral fractionation. The only important differences are: (a) the low- temperature sulfide assemblage is dominated by pyrite (Fig. 11.4b) which is entirely absent in the fractionation case (Fig. 11.3d) and (b) talc appears as a silicate phase in the closed-system case (Fig. 11.4a). Both of these differences are a consequence of the re- dissolution of the chlorite a t low temperature in the closed-system case whereas chlorite is fractionated out of the system (taking Fe and Mg with it) in the fractionation case. Back-reaction of the iron-rich chlorite at 200°C (Fig. 11.4a) provides iron for pyrite, magnesium for talc, and extra aluminum for K-spar.

The purpose in executing the cooling-only calculation (Fig. 11.5) was to isolate the effects of temperature change itself from the effects of boiling. The cooling-only calculation results are valuable for this purpose, but are not necessarily valuable as a model of reality. The single most significant difference between the cooling and boiling calculations is that dissociation of weak acids upon cooling causes pH to decrease whereas pH increases in the boiling calculation (Fig. 11.6). This is why very little sulfide precipitates and phyllosilicates dominate instead of feldspar and carbonate (Fig. 11.5a). The decrease in pH with cooling displaces sulfide into H,S (reaction 6) accounting for the decrease in mHS- in Figure 11.5b. Despite this decrease, metal-bisulfide complexes (Fig. 11.5d) do not dissociate sufficiently to overcome the effect of the acid conditions in enhancing sulfide mineral solubility.

Another conspicuous difference in the cooling calculation is that graphite precipitates. The initial water is enriched in aqueous methane which escapes into the gas upon boiling. If boiling is not allowed, the hydrocarbon is forced to go into graphite upon cooling. (Graphite is the only reduced carbon phase, except methane, available in the numerical model.)

SUPER- AND SUB-ISOENTHALPIC BOILING

In addition to the nearly isoenthalpic boiling and cooling-only cases discussed above, we explored numerically the closed-system geochemical space corresponding to fluid ascent with partial conductive cooling and fluid ascent with excess enthalpy boiling. These results are shown in Figure 11.7 in diagrams of temperature E. gas weight percent. The figures are contoured with mineral stability boundaries, pH, and mole percent C 0 2 gas in the gas phase. The diagrams show the isoenthalpic boiling path (A-D, Fig. 11.7a)

Minerals

Gas Fraction (W. 96) Gas Fraction (wt. %)

co2 Mole Percent

Figure 11.7. Calculated results of closed- system boiling of a Broadlands-like water under a range of conditions corresponding to super-isoenthalpic to conduct ively

0 cooled. The abcissa expresses the gas fraction (including H20 and the dry gases) of the total (liquid plus gas) in weight percent. At any given temperature on the ordinate, only one pressure yields a gas fraction corresponding to isoenthalpic boiling conditions, indicated by curve A-D, in part a). Other pressures yield more or less gas, corresponding to super- and sub- isoenthalpic boiling conditions. See text. a). Mineral phase diagram showing fields for the presence or absence of the indicated minerals. The shaded area at lower left is the stability field for kaolinite; to the right of the shaded area, moscovite is stable rather than kaolinite. See Fig- ure 11.3 for mineral abbreviations. b). pH of boiling waters. Dotted line shows isoenthalpic boiling curve, A-D from part a). c). Mole-percent C02 in the gas phase in equilibrium with the boiling waters. Con- tours for 70 and 90% CO are not labeled

I I I but are shown to the legt of the 50% con- 0 10 20 30

Gas Fraction (wt. YO) tour at low temperature and low-gas fraction. The dotted line shows the isoenthalpic boiling curve, A-D from part a). The other small circles show the points for which calculations were done to produce Figure 11.7.

260 CHAPTER 11

used in t h e previously discussed calculations depic ted in Figure 11.4 and the cooling-only path (A-B, Fig. 11.7a), represent ing complete conductive cooling a s depic ted in Figure 11.5.

The isoenthalpic curve (Fig. 11.7) represents the temperature-s team fract ion t ra jec tory of a fluid which exchanges no h e a t with i t s wall rocks (Elder, 1965; Grant e t al., 1984). All h e a t used t o vaporize t h e water is supplied f rom the wa te r itself. As t h e h e a t is removed, t h e wa te r cools. An isoenthalpic path would be expected under s teady-s ta te flow conditions in a hydrothermal system such t h a t t h e thermal gradient in the wall rock has been previously fixed by t h e boiling fluid (Grant et al., 1984).

All t ra jec tor ies between A-D and A-B (Fig. 11.7) represent fluids t h a t boil during ascent , but which simultaneously lose hea t t o the wall rock. Such t ra jec tor ies apply when pressure increases over t ime due t o self-sealing, such t h a t t he depth of f i rs t boiling ascends (from H toward L, Fig. 11.1). As t h e depth of boiling becomes shallower (because boiling is prohibited by increased pressure due t o self-sealing) ascending hot waters encounter rock t h a t is cooler than boiling along t h e newly established pressure-depth regime allows; thus, the rocks e x t r a c t hea t f rom t h e water. Depending on t h e t i m e r a t e of self-sealing (and consequent r a t e of pressure change with t ime) the e x t e n t of such heat ext ract ion by t h e rocks could resul t in cooling t ra jec tor ies spanning much of the range between A-D and A-B (Fig. 11.7a). Super- isoenthalpic t ra jec tor ies , lying to the right of A-D (e.g., A-E, Fig. 11.7), represent boiling fluids t h a t e x t r a c t excess hea t f rom the wall rock a s they ascend. This produces s team quantit ies in excess of t h a t which t h e ascending water alone could produce. Super-isoenthalpic conditions would prevail, for example, when new hea t is introduced a t shallow levels by magma intrusion. A more common cause would be breaking of a self-sealed system, resulting in a reversion to hydrostatic pressure from super- hydrostatic conditions. This would cause t h e depth of f i rs t boiling t o deepen (downward toward H, and below, Fig. 11.1), exposing hot wall rock t o newly lowered pressure conditions. These s a m e thermal e f f e c t s a r e discussed by Goguel (1982) and Truesdell (1979) in t h e context of exploitation of geothermal energy reservoirs.

The common occurrence of self-sealing and re- breaking in ac t ive geothermal systems is well established (e.g., White e t al., 1971; Muffler e t al., 1971; Keith et al., 1978; F a c c a and Tonani, 1967; Henley and Ellis, 1983) a s is t h e common occurrence of hydrothermal breccias in epi thermal o re deposits (e.g., Berger and Eimon, 1983; Nelson and Giles, 1985; Hedenquist and Henley, 1985). Thus, i t is more probable t h a t many of the sub- and super-isoenthalpic t ra jec tor ies represented in Figure 11.7 a r e visited during the ac t ive l i fe t ime of epithermal ore-forming systems. Chemical implications of this a r e discussed fur ther below.

The results shown in Figure 11.7 a r e based on calculations of overall heterogeneous equilibrium fo r each of t h e individual points shown in Figure 11 .7~ . Most of t hese points were calculated using isothermal pressure-drop t raverses a t various temperatures .

Although each such t raverse is in teres t ing in i t s own right, i t i s most useful t o combine a l l results on a f ew diagrams, then explore t h e various ascent t ra jec tor ies discussed above. The mineral phase boundaries shown in Figure 11.7a a r e bes t understood a s a consequence of t h e pH variations shown in Figure 11.7b, which a r e primarily a consequence of the C 0 2 degassing represented in Figure 11 .7~ . The rapid pH increase with ear ly boiling (corresponding t o small gas fractions, Fig. 11.7b), also apparent in Figures 1 1 . 3 ~ and 11.4b, a r e a result of t he ear ly rapid escape of most C O into t h e gas phase a s indicated in Figure 1 1 . 7 ~ a n 2 plotted in moles in Figure I l.3a. Figure 11.7b shows t h a t essentially a l l boiling paths emana t ing from point A result in pH increases excep t conductive cooling paths near path A-B. The monotonic dec rease in pH along path A-B is shown in Figure 11.5b.

The pH control on gangue s i l ica te mineralogy, for example, is apparent from t h e approximate parallelism of the muscovite-K-spar boundary in Figure 11.7a and t h e pH contours of Figure 11.7b. K-spar forms on t h e high-pH side of the boundary in accordance with reaction 8. According t o the phase diagram, a l l boiling t ra jec tor ies t h a t result in more than 10 percent gas a t 1 0 0 ~ ~ will produce deep muscovite and shallower K-spar. The particular depth where t h e transit ion f rom muscovite t o K-spar occurs i s f ixed by the point of in tersect ion of the ascent t ra jec tory (e.g., A-C) with t h e phase boundary in Figure 11.7a.

The cinnabar field in Figure 11.7 is particularly ins t ruct ive because i t demonstra tes t h e e f f e c t s of compet ing react ions on t h e solubility of a sulfide mineral. According t o Figure 11.7a, isoenthalpic boiling does not precipi ta te cinnabar f rom this water, but neither does conductive cooling. Consider an isothermal P-drop t raverse ( lef t t o right, Fig. 11.7) a t a t empera tu re between 1 0 0 ~ ~ and 170°C; cinnabar remains soluble in the absence of boiling (zero gas f rac t ion, Fig. 11.7) because t h e large H S concentra t ion in the aqueous phase displaces t z e following reaction t o t h e l e f t

HgS(H2S)2 = HgS + W2S ( 9 ) ( a q ) ( c i n n a b a r ) ( a q )

The protonated mercury complex and H2S a r e strongly favored by the low pH (Fig. 11.7). Upon boiling, react ion 9 is displaced t o the right a s H S (gas) escapes, resulting in cinnabar precipitation. dowever , a s boiling proceeds, t he consequent increase in pH (Fig. 11.7b) shifts sulfide equilibria such t h a t t h e mercury-sulfide complex, HgST, increases in concentration, causing cinnabar to redissolve ( react ion 10) despite t h e fur ther loss of H2S gas

HgS + HS--+ HgS= + H+ (10 ) ( c i n n a b a r ) ( a q ) (aq? ( a q )

In t h e con tex t of possible fluid ascent t ra jec tor ies , this behavior of cinnabar means t h a t only paths with par t ia l conductive cooling of the fluid will produce cinnabar.

The cinnabar and muscovite-K-spar examples i l lus t ra te t h e profound e f f e c t of t he boiling hea t budget on t h e expected epi thermal vein mineralogy. T o the e x t e n t t h a t t h e hea t budget of boiling fluid "parcels" varies through t ime, fluid t ra jec tor ies on Figure 11.7a vary and vein mineralogy must vary. This is one probable source of mineral banding in epi thermal veins. As a system experiences self-sealing and re-breaking, t ra jec tor ies shi f t from isoenthalpic t o sub-isoenthalpic to super-isoenthalpic on Figure 11.7a. Such a fluid, a t i t s 230°C isotherm (which itself shi f ts in space) would f i rs t precipi ta te quartz, ca lc i te , K-spar, Fe-chlorite, chalcopyrite, sphalerite, and galena (Fig. 11.4a), then (depending on where t h e sub- isoenthalpic t ra jec tory falls) perhaps quartz, muscovite, pyrite, and sphalerite (but not galena, chalcopyrite, chlorite, K-spar, and calcite), a s indicated by Figure 11.7a a t 230' and 5-percent gas fraction. A t th is s t age t h e much lower pH of the fluid (Fig. 11.7b) might also e t c h ear l ier formed minerals and cause muscovite t o replace ear l ier formed K-spar (e.g., Hedenquist and Henley, in preparation, 1985). Upon re-breaking and a swing t o super-isoenthalpic conditions, the fluid again would precipi ta te carbonate, feldspar, etc., a t i t s 230' isotherm.

Instead of following a fluid parcel of given t empera tu re through t h e history outlined above, we may consider how physical and chemical conditions vary a t a point fixed in ,space (in a vein cavity, for example). A gradual pressure increase due t o sealing would correspond to traversing toward high t empera tu re in Figures 11.3 and 11.4 for any given position in a vein. Thus, a point initially a t 250°C (Fig. 11.3b) where quar tz , K-spar, chlorite, sphalerite, and galena a r e precipitating could experience a shi f t t o 2 7 8 ' ~ a s pressure increases, result ing in e tching of t h e previously formed minerals, overlapping precipitation of muscovite, and replacement of K-spar by muscovite. Subsequent pressure drop would result in a re turn t o precipitation of the lower t empera tu re assemblage. Other more complicated changes could occur, depending on t h e r a t e of pressure change, result ing in sub- and super-isoenthalpic boiling (Fig. 11.7a) a t t h e given point.

BOILING AND GOLD PRECIPITATION

The large sulfide content of the Broadlands-like water used here stabil izes t h e Au(HS)Z complex (Fig. 11.3h) precluding gold precipitation throughout t h e boiling range from 2 7 8 ' ~ to 1 0 0 ~ ~ . In t h e calculation of boiling with mineral f rac t ionat ion (Fig. 11.31, gold is undersaturated by more than two orders of magnitude a t 2 7 8 ' ~ (log Q/K = -2.44), but th is decreases t o undersaturation by half an order of magnitude a f t e r boiling t o 1 0 0 ~ ~ (log Q/K = -0.57). Thus, if t h e solution had been sa tura ted with gold a t t he point of incipient boiling (as in Drummond and Ohmoto, 1985, who used a large gold concentration) or if we had used the new value for gold concentra t ion from Henley and Brown (1985, this volume) which is 2 3 t imes g rea te r t han the 0.1 ppb t h a t we assumed (Table 11.1), gold would have precipitated due t o boiling. The

c r i t i ca l react ion for gold precipitation in the ac id t o neutra l pH range is t h e following

8Au (HS) 5 + 6H+ + 4 ~ 2 0 ~ A U O + S q + I 5H2S ( a q ) ( a q ) ( a q ) ( g o l d ) ( aq ) ( g o r a q )

( 1 1 )

This reaction enta i ls reduction of Au+ t o AuO by sulfide which is thereby oxidized t o sulfate. React ion 11 makes i t c lear t h a t pH increase induced by boiling competes with loss of H2S gas in determining whether gold precipi ta tes or not (see also, Drummond and Ohmoto, 1985). The thermochemical d a t a used here for gold-bisulfide complexes (Seward, 1973) indicate t h a t a s pH increases in t h e range calcula ted he re (6.4 t o 7.81, gold solubility increases (reaction 11) a s H2S dissociates t o HS-. Thus, the ca lcula ted approach toward gold sa tura t ion upon boiling is principally a consequence of loss of H2S t o the gas phase, displacing react ion I I toward gold saturation (see also, Hedenquist and Henley, 1985).

The precipitation of sulfide minerals const i tu tes another sink for sulfide which promotes gold precipitation. However, in the case calculated, t h e e scape of H2S to t h e gas phase is f a r more significant than sulfide mineral format ion a s a sink for aqueous sulfide (compare Figs. 11.3a and 11.4a), a s discussed in a preceding section. In any case, gold did not precipi ta te because of t he excess sulfide content of original water. Instead, i t s tayed in t h e aqueous phase and, in t h e geologic context , i t would be t ranspor ted to t h e surface hot-spring environment (see below).

In order fo r gold to precipi ta te with sulfides in t h e deep vein environment a s i t did, for example, a t Comstock (Bastin, 19221, i t is necessary t o c r e a t e a sulfide-deficient water by removing essentially a l l of t h e aqueous sulfide t o minerals and gas. This is best accomplished by boiling waters which have a large r a t io of me ta l s (Fe, Cu, Pb, Zn) t o sulfide so t h a t precipitation of pyrite, chalcopyrite, galena, and sphaler i te along with H2S gas removes a l l sulfide f rom t h e aqueous phase. The low salinity of t h e Broadlands- l ike water precludes high-metal contents because chloride is necessary to complex significant concentra t ions of base meta ls in the presence of high- sulfide concentrations. Other calculations (Reed, 1985) on more saline waters (e.g., .5 m NaCI) show t h a t chloride complexes of Pb, Zn, and C u dominate, even in high-sulfide waters. A large metal/sulfide r a t io is necessary t o assure sulfide depletion with consequent gold precipitation, and large salinit ies a r e necessary t o large base-metal concentrations. Thus, t h e boiling of saline waters should produce gold precipitation with sulfides in accordance with t h e following example react ion in which galena and gold precipitation a r e coupled

4 8 ~ ~ ' + so4 + 30 CI - + 15PbS + 24H' (12) ( g o l d ) ( aq ) ( aq ) ( g a l e n a ) ( a q )

As boiling proceeds, COZ escapes, driving up pH (reaction 1); th is decrease In H+ act iv i ty drives sulfide

precipitation (reaction 31, which in sulfide-deficient waters is coupled t o gold precipitation a s in reaction (12). Thus, t h e format ion of a hot-spring gold deposit (see below) a s opposed to a base-metal-vein gold deposit may result primarily f rom the distinction between sulfide-deficient a n d sulfide-excess boiling waters, which itself i s t i ed t o the distinction of low- salinity = high-salinity waters. Fluid-inclusion d a t a support this distinction.

THE HOT-SPRING ENVIRONMENT

In the hot-spring environment a t and just below t h e surface over a boiling hydrothermal system, t h e e f f ec t s of cool temperatures , a tmospher ic oxidation, and t h e influx of me teo r i c ground wa te r s determine pat terns of rock a l tera t ion, water chemistry, and o r e formation. Figures 11.1 and 11.2 show t h e near- surface geochemical processes t h a t we chose t o simulate numerically for t h e purpose of understanding the origins of hot-spring gold deposits (Berger and Eimon, 1983; Henley and Ellis, 1983). In t h e process of unravelling t h e chemistry of gold deposition, we generated numerical models of t he format ion of acid condensates, acid-sulfate waters, and neutra l carbonate-sulfate waters.

All results were produced using the same fundamenta l approach for multi-component heterogeneous equilibrium calculations (Reed, 1982) t h a t we applied to boiling. The sequence of operations on t h e gas and liquid phases l e f t over from boiling a r e a s follows (refer t o Figs. 11.1 and 11.2 for points designated below by letters): 1. condense the gas phase in open space by cooling

from IOOOC to 93OC (B). 2. Oxidize the gas with a tmospher ic 0 2 (C). 3. Condense the gas in cold, oxygenated ground

water and r e a c t a t 99OC (D). 4. Ti t r a t e (D) the acid-sulfate wa te r f rom (C) in to

the boiled aqueous phase from (A). 5. T i t r a t e (F) t h e cold, oxygenated ground water in to

the boiled aqueous phase f rom (A). The results of e a c h of these calculations a r e presented below in t h e order of t he above list.

Condensation of t h e Boiled Gas

When t h e gas phase from a boiling hydrothermal system escapes into cool, f r ac tu red rock above t h e boiling water table, t h e H 2 0 f ract ion largely condenses (White, 1957; White e t al., 1971) and carr ies with i t small amounts of dissolved C 0 2 and H2S. Subsequent, d i rec t a tmospher ic oxidation (Fig. 11.8) produces acid-sulfate waters. We calculated t h e chemical character is t ics of the condensate in order t o provide a s tar t ing composition fo r oxidation, but also t o understand be t t e r how the condensate may a c t alone in a l ter ing wall rocks and re-dissolving vein minerals.

We separa ted the gas phase f rom the li uid phase 73 l e f t over from the boiling calculation a t 101 C (Table

11.1, Fig. 11.3a) and numerically condensed i t by cooling the gas alone t o 93OC a t a pressure of 1.01 bar. Computationally, such a condensation is no

di f ferent than t h e boiling calculation. There is a gas phase in equilibrium with a liquid phase. As t empera tu re decreases at constant pressure, gas condenses t o liquid and chemical potentials of all species a r e equalized between t h e two phases. The resultant liquid composition at 95OC is given in Table 11.2 and pH and gas composition a r e shown in Figure 11.9. Gas and liquid compositions a r e also shown a t t h e ex t r eme left-hand ends of Figures 11.10a and 11.10b.

Atmospheric 0 2

T°C Conden- sate

100

(H20,CO2,H2S) * 0 2 Tqi

H20 C02 Liquid Condensate

95OC

B

Figure 11.8. a). Schematic diagram showing the boiled gas phase passing through a temperature gradient in fractured rock above the water table. The H20 condenses, carrying some C02 and H2S with it. Atmospheric oxygen enters from above, oxidizes HZS to sulfuric acid which dissolves in the liquid condensate. b). Diagram of a semi-closed system approximation to the scheme in part a), as set up for numerical oxidation at 9 5 O ~ (see text and Fig. 11.10). Oxygen gas is titrated into a two-phase mixture of condensate plus gas, but the original gas phase remains in contact with the liquid throughout the calculation. Thus, the H2S is gradually consumed by oxidation, rather than being continuously re-supplied from belm.

M. H. REED & N. F. SPYCHER 263

Table 11.2--Water compositions

Cold ground water

PH 7 (25'~) (molality) (ppm)

C1- .733e-04 2.6

F- .174e-04 0.33

C O ~ .202e-02 121

HS- --- --- CH4 aq. --- ---

S O ~ .323e-04 3.1

O2 aq. .250e-03 8

Si02 .105e-02 63

~ 1 + + + .l lle-06 0.003

~ a + + .724e-04 2.9

M ~ + + .535e-04 1.3

~ e + + .179e-06 0.01

K + .358e-04 1.4

~ a + .457e-04 10.5

Condensed Acid-sulfate water Gas + ground water gas (95'~) (0.03 moles of added 02) (72 kg of added water)

5 0 : . . 1 ~ ~ . ~ . - - 1 ~ - - .

40 :

30 7

20 :

1 Condensation 1 bar, Oar Phars

t . L . 1 . . 3 I . . . 1 . . . 1

: Condensate Iliquid)

4.0% 92 94 96 98 18%

Temperature <Deg.C.)

Figure 11.9. Calculated process of condensation of t h e gas phase boiled o f f the Broadlands- l i k e water. The C02 ,con ten t of t h e gas phase i n equilibrium wlth the l iqu id during condensation from 1 0 0 ~ ~ to 93O~ a t 1.01 bar is g iven i n a ) . The l i q u i d c o n t a i n s wa te r with dissolved H2S, a2, etc., yielding the pH shown i n b). pH decreases with cooling a s more and more C02 d i s s o l v e s i n t h e aqueous phase, d e s p i t e t h e f a c t t h a t t h e p r o p o r t i o n of C02 i n t h e gas phase in- c r e a s e s a) .

During condensation, p,H decreases from 4.9 at IOOOC to 4.5 a t 9 5 ' ~ (Fig. 11.9) while the CO fraction of the gas phase increases from less than 18 percent to more than 40 percent by weight (2.4 to 18 mole percent). The remainder of the gas is H20, except for CH4, H2, and H2S, which also increase in proportion, but which constitute less than 1 percent of the gas. The pH of 4.5 at 95OC is entirely a consequence of carbonic acid, the molality of which increases during cooling (as the solubility of C02 increases) despite the fact that the proportion of C02 in the gas phase also increases. Condensation thus produces a mildly acidic liquid, which a t 9 5 ' ~ has a mass of 353 grams per initial 1.015 kg of unboiled water. The associated gas is quite C 0 2 rich, as also indicated at low temperature and small gas fraction in Figure 1 1 . 7 ~ where CO;, gas fractions exceed 70 mole percent. Such CO - r~ch gases reported from the boundaries of hot hyirothermal systems (Mahon e t al., 1980; Coguel, 1982) are undoubtedly a consequence of condensation of the H20 gas fraction as calculated here.

The calculated pH of 4.5 for the condensed liquid is certainly acidic enough and its volume great enough to effect alteration of volcanic wall rocks to illite or kaolinite (depending on temperature and K+ activity), and i t may be responsible for some of the illitic wallrock alteration reported in the upper parts of vein systems, particularly in the hanging wall (Berger and

Oxidation of gases 1

1 CI-

Oxidation of gas

Moles of oxygen added

Eimon, 1983). The condensate is not ac idic enough, however, t o cause t h e extensive a l tera t ion of rocks a t shallow levels t o kaolinite and alunite. For this, acid- sul fa te wa te r s a r e necessary.

Oxidation of Cases t o Produce Acid-Sulfate Waters

Acid-sulfate waters of very low pH a r e produced by d i r ec t a tmospher ic oxidation of sulfidic gases above t h e wa te r t ab le (White, 1957) a s indicated in Figures 11.1 and 11.2. In a following section, i t is shown t h a t oxidation of gases by dissolved oxygen in cold ground water cannot produce acid-sulfate waters because t o o much concurrent dilution precludes a low pH. The physical concept of t he oxidation calculation is i l lustrated in Figure 11.3a, which shows t h e gas condensing in a f r ac tu re a s i t t raverses a t empera tu re gradient. The liquid condensate remains in con tac t with excess (non-condensible) gas which continues t o flow through the f r ac tu re while a tmospher ic oxy e n Pb gains access from above. The oxygen r eac t s a t 95 C with t h e two-phase mixture (liquid plus gas, Fig. 11.8b) remaining a f t e r condensation (i.e., we do not exclude t h e remaining non-condensible gases from react ing with oxygen along with the consti tuents dissolved in t h e liquid). The calculation is e f f ec t ed by numerically t i t r a t ing O2 gas in to t h e condensate plus gas a t 9 5 ' ~ and 1.01 bar with computation of overall heterogeneous equilibrium a f t e r each t i t ra t ion step.

From t h e standpoint of numerical computation, carrying ou t a continuous oxygen t i t ra t ion react ion such a s th is one (Fig. 11.10) and others described below enta i ls overcoming a significant obstacle a t t he transit ion f rom reduced t o oxidized conditions (e.g., .0145 moles 02, Fig. 11.10). A t this point, t he

Figure 11.10. Calculated oxidation of 353 grams of the boiled gas phase by atmospheric O2 to produce an acid-sulf ate water. a). The composition of the gas phase in equilibrium with the aqueous phase as 0 is titrated into the liquid p$us gas mixture. CH9 was allowed to oxidize despite the fact that such oxidation is kinetically retar- ded. See text. b). Composi- tion of the aqueous phase, showing molalities of individual species. Activity of hydrogen ion is also shown, for which the ordinate refers to log activity.

concentra t ions of many reduced sulfide species (HS-, H2S, and complexes involving these) become so small t h a t thei r molali t ies cannot be reliably computed using normal machine double precision (64-bit representa t ion of variables). A t t h e same point, concentrations of aqueous 0 2 and H 0 became large. Under such oxidizing condition: $e really do not c a r e t o know about sulfide molali t ies less than so i t is convenient t o recas t all redox equilibria in t e r m s of t h e H 2 0 - 0 2 redox pair in place of the HS--SO= ai r used under reducing conditions (Reed, 1982). !his amounts t o substi tuting 0 2 for HS- a s a component species, then providing a whole d i f ferent s e t of equilibrium constants for t he various redox equilibria. We a r e able t o overlap t h e computations slightly a t t he transit ion point t o verify t h a t a l l computed molali t ies m a t c h through t h e transition.

Figure 11.10b shows t h a t a f t e r react ing less than .02 moles of O2 with 353 grams of condensate plus gas, t h e pH changes from 4.5 t o 2.06 a s sulfuric acid is produced from H2S

H2S + 202 -) H SO (13) ( g a n d a q ) ( g ) faq!

( s u l f u r i c a c i d )

The overall process occurs in t w o steps. The f i rs t is oxidation of me thane and hydrogen (Fig. 11.10a), which a r e fully oxidized by nearly .O1 mole of 02 . In nature , th is oxidation is almost certainly insignificant because of kinetic barr iers t o oxidation of reduced carbon and hydrogen a t low t empera tu re (see review of this question by Drummond and Ohmoto, 1985). We allowed this oxidation because i t has no quant i ta t ive

M. H. REED & N. F. SPYCHER 265

e f f e c t on results e x c e p t t o delay t h e oxidation of H S. While me thane and hydrogen oxidize (Fig. I ?.10a), pH remains constant (Fig. 1 l.lOb) and sul fa tes begin a gradual increase in molality. Once t h e hydrocarbon oxidation is complete, oxygen a t t a c k s H2S, producing sulfuric acid and a ca tas t rophic increase in acidity t o a pH of 2.06 with consequent shi f ts in sulfide, carbonate , and su l f a t e equilibria a s shown in Figure 11.10b. As HZS oxidation proceeds, t h e oxidation products dissolve In t h e aqueous phase, deple t ing H2S from t h e gas (Fig. 11.10a). (The calculation Itself provides no information a s t o whether t h e ac tua l oxidation of H2S occurs in t h e gas o r aqueous phase, o r a t their interface.) The final disappearance of H2S gas ( a t .0145 moles 0 2 ) allows O2 gas t o accumula te for t he f i r s t t ime in t h e gas phase. A resul tant water composition is given in Table 11.2. The particular pH a t t a ined depends on how much H S is oxidized. If only t h e H.2S origi ally d issolvebin the aqueous phase were oxidlzed (lo1 m, Fig. 11.10b), t h e pH would not have dropped below 3.4. Since many acid-sulfate wa te r s have pH's between 2 and 3 and some a r e less than 2 (White et al., 1971; Rowe e t al., 19731, i t appears t h a t oxidation of H2S gas t h a t did no t originally dissolve in t h e condensate contr ibutes t o the sulfuric acid in natura l acid-sulfate springs. This conclusion applies in general, even though i t i s based he re on a single, speci f ic water plus gas composition, because our Broadlands-like colnposition is unusually rich in sulfide gas and represents an e x t r e m e case.

Clearly such a n acidic water would quickly a t t a c k i t s host rock producing kaolinite and leachi base cations. Among such dissolved consti tuents, A1 38 i s significant with respect t o the gangue in gold ores. This point i s discussed below in connection with mixing of acid-sulfate waters with the boiled liquid (process E, Figs. 11.1 and 11.2), but f i r s t we discuss gas condensation and oxidation in me teo r i c ground water.

React ion of Cases with Meteoric Ground Water

In order t o evaluate the potential of oxygenated ground water for producing acid waters t o r e a c t with gold-bearing boiled waters, we t i t r a t ed a New Zealand me teo r i c ground water (Table 11.2, col. I , Timperley,

1983) containing 8 pprn dissolved O2 in to t h e gas phase l e f t over f rom boiling a t 10l°C and 1 bar. This produces a neutra l bicarbonate-sulfate water (Fig. 11.11) of t h e so r t discussed by White et al. (1971) which they postulated formed by just such a reaction. As shown in Figure 11.1 1, t h e basic pa t t e rn of change in sulfide and sul fa te is t h e s a m e a s in Figure 11.10, excep t t h a t t h e process is "s t re tched out" because 60 kg of p o u n d water containin 8 ppm 0 2 (2.5 x 10- moles/kg H20 , Table 11.27 a r e required t o oxidize a l l C H , H2, and H2S in t h e 353 grams of gas from boiling. Bigure 11.1 I shows a gradual increase in aqueous SO$ as H2S oxidizes, finally resulting in a sulfate-carbonate water with a pH of 6.3 a t 99OC (Table 11.2, Fig. 11.11). The gas phase changes composition in this react ion in exact ly t h e s a m e way a s shown in Figure 11.10a, excep t t h a t t h e positions of slope change a r e "s t re tched out" re la t ive t o Figure 11.10, just a s in t h e aqueous phase (Fig. 11.1 1).

In t h e ground wa te r t i t ra t ion reaction, pyrite forms continuously between 0 and 58 kg of t i t r a t ed water and is replaced by hemat i t e a t 58 kg (Fig. 11.11) where O2 becomes abundant in t h e aqueous phase. The pyr i te forms a s iron, supplied by t h e cold ground water , and r eac t s with H2S supplied by t h e boiling hydrothermal water

4 ~ e ~ + + 7H2S + D4 -+ 4FeS2 + 4 H p + 6H+ (14) ( aq , (gas ( a q ) ( p y r ~ t e ) ( a q ) ( a q ) g.w.) o r a q )

To t h e ex ten t t h a t inflowing ground wa te r s ca r ry significant concentrations of iron, this process is likely t o be important in producing a "pyrit ic halo" on epi thermal o re deposits a s indicated in Figure 11.12. (Furthermore, i t would ope ra t e equally well t o produce t h e well-known pyr i t ic halos on porphyry copper deposits.) Kraynov e t al. (1982) show t h a t iron concentrations in cont inenta l ground wa te r s commonly range from a few pprn t o severa l t ens of ppm. Such concentrations, when r eac ted with H2S, a r e sufficient t o produce a n excess of pyr i te (mostly in veins) beyond t h a t which is produced by sulfidation of iron originally in the host rock itself.

Figure 11.11. Calculated composition of the aqueous phase resulting from mixing and reaction of oxygenated ground water with the condensate plus gas from the boiled water, producing a neutral carbonate-sulfate water. Pyrite precipitates in the reduced interval (0-58 kg) and hematite precipitates in the oxidized interval (above 5 8 kg) .

K g . of water added

Ground Water

I Mixing and Fiea&on Ppt. Gold and Kaolinite

1 Ascending Boiling Waters with Sulfide Excess and Gold

Gold Precipitation f rom Mixing of Acid-Sulfate Water with Boiled Aqueous Phase

In a preceding sect ion on boiling and gold precipitation, we pointed out t h a t boiling does not cause gold t o precipi ta te from a sulfide-excess water with low gold concentration. Such waters carry their gold t o t h e near-surface hot-spring environment where gold may precipi ta te when the Au(HS)Z complex is destroyed by oxidation of the sulfide. Another reaction, possibly of g rea t e r significance than oxidation itself t o near-surface gold precipitation, is acidification of gold-bearing boiled waters by descending acid-sulfate wa te r s (process E, Figs. 11.1 and 11.2; Fig. 11.12). In this calculation, we numerically t i t r a t e a t 95OC the acid-sulfate water produced by oxidation of t h e boiled gas (see above and Fig. 11.10) into the liquid remaining a f t e r boiling of our original water with mineral fractionation (Fig. 11.3). Results a r e shown in Figure 11.13. A s teady increase in H+ act iv i ty a s acid water is added is apparent in Figure 11.13b. This dominates a l l chemical reactions. Acidification causes cinnabar to precipi ta te a s react ion 10 (see above) is displaced t o t h e left . Af t e r approximately 0.35 kg of ac id water is added t o t h e original 0.650 kg of boiled water , a gas phase separa tes (Fig. 11.13a) a s C 0 2 and H2S a r e produced by reactions 1 and 2 (see above). Analogous reactions for HF and HCI a r e also displaced toward the associated form by acidification. Because the gas phase is necessarily a solution, other gases join H S, C 0 2 , HF, and HCI a s shown in Figure l l . l $a . Separation of H2S gas combines with t h e acidification t o break up the dominant complex of copper, Cu(HS)j, forcing precipitation of covelli te (Fig. 1 1 . 1 3 ~ )

Figure 11.12. Schematic diagram of fluid-mixing zones in the upper part of a biling hydrothermal system. In the center of the diagram, ascending, boiled, gold-bearing waters (as in Figs. 11.3 and 11.4 and Table 11.1) encounter descending acid-sulfate waters produced above the water table (Figs. 11.8 and 11.10). The waters mix and react precipitating gold (Fig. 11.13). The descending waters also alter the wall rock to kaolinite, particularly in the hanging wall of the vein. The right-hand side of the diagram shows mixing and reaction of biled gases (with cold, recharge ground water), producing a neutral carbonate- sulfate water (Fig. 11.11). If the recharge water is iron rich, reaction of H2S gas with iron could produce pyrite, which would appear as a "pyritic halo" on the ore deposit.

l o p + 8Cu(HS)Z + SO: ( a q ) ( a q ) ( a q )

--)8CuS + 4H20 + 9H2S ( 1 5 ) (cove1 l i t e ) ( a q ) ( g a s and a q )

Fur ther acidification similarly destroys the dominant gold complex, Au(HS)Z, forcing gold precipitation

6Hf + 8Au(HS)2 + 4~20--+8AuO + + 15H2S (16) ( a q ) ( a q ) ( a q ) ( g o l d ) ( a q ) ( a q and

g a s )

The deta i led numerical results show t h a t t h e dominant e f f e c t in both of these reactions i s displacement of bisulfide in to aqueous H2S by acidification (note declining HS-, Fig. 11.13b), not t he escape of sulfide t o t h e gas phase, although t h e l a t t e r does remove 10 percent of t he original aqueous bisulfide.

For this calculation, we used numerical acid- sul fa te water (Table 11.21, which had not yet r eac ted with i t s numerical host rock. Cons quently our acid- su l f a t e wa te r lacked significant Alg', in contras t t o the high concentra t ions measured in natural acid- su l f a t e waters (e.g., 4 t o 30 ppm a t Yellowstone, Rowe e t al., 1973). If A I ~ ' were present in the ac id wa te r when i t mixed with t h e hot, ascending boiled water , kaolinite would b e expec ted t o precipi ta te

2 ~ 1 ~ + + 2 s i 9 ( a q , descend ing) ( aq , a scend ing)+ :$

+ A ; z S i P S ! M ) 4 + 6I-P (17 ) a o 1n1 t e ) ( aq )

Such kaolinite would precipi ta te along with gold a s

M. H. REED & N. F. SPYCHER

Figure 11.13. C a l c u l a t e d r e a c t i o n and p rec ip i t a t ion of gold upon mixing of a c i d - s u l f a t e wa te r w i t h t h e l i q u i d remaining a t 1 0 0 ~ ~ and 1 bar from boi l ing of t h e Broadlands-like water. a). Composit ion of t h e gas phase which s e p a r a t e s a s a consequence of a c i d i f i c a t i o n . b). M o l a l i t i e s of i n d i v i d u a l spe- c i e s i n t h e aqueous phase . A c t i v i t y of hydrogen ion is a l s o shown. c) . Ploles of min- e r a l s t o p r e c i p i t a t e from 650 grams of boi led liquid.

I " ' I " ' I " ' "20 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Liquid + acid Gases

-No Gas Separated-

I r - 7 , , I _____-_--- - -

Liquid + acid

vein gangue (Fig. 11.12) where the waters mix. An intimate mixture of gold with pure white kaolinite between bands of vein quartz has been reported by Kurt Katsura (U. Oregon, personal communication, 1985) in upper parts of veins in the Champion Mine, Bohemia District, Oregon. Kaolinite with gold in other such occurrences elsewhere might be mistaken for alteration clay "washed in" from altered wall rock.

Gold is commonly associated with arsenic and antimony sulfides around certain hot springs and in some epithermal ores (White, 1981; Henley, 1985, this volume.) Henley (1985, this volume) suggests that the As and Sb sulfides precipitate owing to acidification, and Ellis (1969) has suggested that the colloidal precipitating sulfides scavenge gold. Our calculations indicate that gold, too, precipitates by acidification (reaction 16); thus, the association of arsenic-antimony and gold (and cinnabar, Fig. 11.13~) may actually be a consequence of co-precipitation of all from sulfide complexes by acidification.

I " ' I ' ' ' I ' ~ ~ _

Liquid + acid - Minerals - -

cv ,-------------------------- - cinn - -_------ , gold -

/ -

r' , , , , , , , i I

-

K g . of acid water added

Gold Precipitation from Mixing of Oxygenated Ground Water with Boiled Aqueous Phase

When gold precipitates by acidification (as above, Fig. 11.13), atmospheric oxidation is indirectly responsible because oxidation of H2S produces the requisite sulfuric acid. The critical gold-precipitating reaction (161, however, is driven by acid, not by oxygen. Another reaction that is capable of driving gold precipitation is oxidation of sulfide by oxygen dissolved in cold ground water. In this section, we discuss such a reaction coupled simultaneously to the cooling of boiled water by the cold ground water.

This calculation is accomplished by numerically titrating (Fig. 11.2) the 25OC New Zealand meteoric ground water containing 8 ppm dissolved O2 (Table 11.2) into the 1 0 0 ~ ~ gold-bearing water remalnlng from boiling (Fig. 11.3). In addition to mass accounting, the enthalpy and consequent temperature of the mixture is calculated from the masses and

known enthalpies of t h e mixing fluids a s discussed by Reed (1985). Thus, a s cold water i s t i t r a t ed in to hot, t empera tu re decreases f rom IOOOC toward 25OC. The results of this process a r e shown in Figure 11.14, using t empera tu re on the abcissa. These graphs a r e bes t r ead f rom right t o l e f t , ref lec t ing progressive addition of cold wa te r t o hot water.

There a r e two s t ages of react ion appa ren t in Figure 11.14, corresponding t o reduced, then oxidized, conditions a s also found in o the r oxidation calculations ( F P . 11.5, 11.10, and 11.1 1). The transit ion occurs a t 65 C (Fig. 11.14), a t which point a sufficient amount of dissolved oxygen had been added f rom t h e cold water t o oxidize t o su l f a t e a l l sulfides originally in the boiled water. A t t empera tu res above 65OC, minuscule quant i t ies of sphalerite, galena, acanthi te , and chalcopyr i te precipi ta te a s t h e cold ground wa te r chills t h e mixture, wringing a l l residual base me ta l f rom t h e boiled water. These sulfides precipi ta te despite dilution because cooling by 35OC (a t constant pH, Fig. 11.14b) reduces thei r solubilities by one or more orders of magnitude while concurrent dilution only reduces me ta l concentra t ions t o less than half. This s a m e e f f e c t would cause sulfides t o precipi ta te by cold water dilution a t high temperatures , too, and may be an important cause of base-metal precipitation in epi thermal veins without boiling. Pyr i t e precipi ta tes abundantly (Fig. 11.14a) a s iron f rom t h e ground wa te r r eac t s with sulfide in t h e hot water. A t just above 70°C, gold precipitates because aqueous sulfide concentra t ion has been sufficiently decreased by oxidation t o destroy t h e gold-bisulfide complex according t o t h e following reaction

8Au(HS)i + 30 4 + 4720 ( a s ) (aq aq )

-+ m u 0 + 1 6 w 4 + 24I? (18) ( g o l d ) ( a q ) ( a q )

Fur ther oxidation of sulfide t o sul fa te resul ts in dissolution of all sulfides a t 6 5 ' ~ a f t e r t h e addition of .56 kg of cold wa te r t o a n init ial .65 kg of boiled liquid.

Figure 11.14a does not show t h e spectacular mineral replacement sequence t h a t is compressed in to a f rac t ion of a deg ree of t empera tu re change around 6 5 ' ~ where log fO2 changes f rom -55 t o -4. In this interval (where pyr i te dissolves and hemat i t e precipitates, Fig. l l . l 4a ) , t h e following reactions occur: (a) chalcopyrite is replaced by bornite, then bornite by chalcocite, ha lcoci te by nat ive copper, and copper by aqueous CuL; (b) acan th i t e is replaced by nat ive silver, then silver dissolves a s aqueous AgCl and AgC12 ; and (c) cinnabar is replaced by nat ive mercury, then mercury dissolves in aqueous chloride and hydrox- ide complexes. Clearly these resul ts a r e applicable t o understanding t h e weather ing of sulfide o re deposits, but th is goes beyond t h e scope of t h e present study.

I t i s apparent in Figure 11.14b t h a t t h e to t a l molali t ies of copper, si lver, and mercury decrease a s pyrife! chalcopyrite, cinnabar, and acan th i t e preclpl ta te owing t o cooling and destruction of thei r sulfide complexes by oxidation. Near 65OC, just before complete oxidation of t h e sulfides, each of these molali t ies reaches a minimum a s the meta ls a r e s tored in sulfide minera ls and nat ive metals. Upon complete oxidation, t hese solids redissolve a s sulfides

I " ~ ~ ~ ' " ' l r r " " " ' l r r " " " ' - - liquid + g . w . -

- -

Figure 11.14. Calculated reaction and temperature change upon mixing of the 1 0 0 ~ ~ boiled water with a 2 5 O ~ oxygenated ground water. a). Moles of minerals to precipitate from 650 grams of initial boiled liquid. Oxygen fuga ch ges abruptly from 10- f2 kz 102 at the point where all sulfide is oxidized to sulfate. In that minuscule interval, additional minerals precipitate as indicated and discussed in text. See Figure 11.3 for mineral abbreviations. b). Total molalities of ore metals in the aqueous phase. The sharp down- turn at low temperature is a consequence of the extreme dilution by cold ground water.

40 60 80 100

Temper at ur e

a r e oxidized and t h e m e t a l s a r e complexed by chloride. The molali t ies of copper, silver, and mercury shi f t back t o maximum values which, because of t h e dilution, a r e no t a s high a s t h e s t a r t i ng values (before mixing). Fur ther dec reases in m e t a l molali t ies occur a s dilution proceeds. The lead and z inc molali t ies show a slightly d i f f e r en t t r end because galena and sphaler i te redissolve before oxidation is complete. Although gold is st i l l present a f t e r oxidation of t h e sulfides, t h e molali ty of gold follows t h e s a m e t rend a s copper, silver, and mercury because gold is car r ied in a bisulfide complex. The gold molali ty a lso drops sharply a t 6 5 ' ~ , be fo re shift ing back t o a higher value. The higher gold molali ty results when, in t h e absence of sulfide, na t ive gold is oxidized in t h e presence of C1- t o a gold chloride complex according t o t h e following react ion

This explains why t h e amoun t of na t ive gold is slightly lower a f t e r oxidation (Fig. 11.14a). Eventually, fu r the r oxidation would redissolve all t h e gold in to AuC14 and AU+++; however, in our case, t h e concentra t ions of au r i c species were insignificant throughout t h e en t i r e mixing process.

Upon comple t e oxidation, t h e only remaining minerals a r e hema t i t e and gold. A t t h e low t empera tu re s involved here , a n assemblage of goethi te and gold would probably prevail in t h e natura l environment. Thus, t h e final product of t h e mixing and oxidation is a n assemblage of gold and iron oxide, with a l l o ther o r e me ta l s dissolved out. This may occur in a n ac t ive hot-spring environment, a s suggested in Figure 11.1, producing a n assemblage t h a t may be mis taken fo r one formed by weathering.

SUMMARY

Using calculations of simultaneous, multi-phase chemical equilibria, we have explored some of t h e const ra in ts t h a t local thermodynamic equilibrium places on t h e mineral con ten t and m e t a l zoning in epi thermal o r e deposits. We focused on a dilute, gas- r ich Broadlands-type water , but many of our conclusions c a n be applied t o o the r waters. A summary of t h e principal conclusions re levant t o epi thermal o r e genesis follows.

Boiling, with consequent re lease of C 0 2 gas, causes pH t o increase by more than o n e unit. This drives precipitation of base-metal sulfides and carbonates. Most of t h e C 0 2 originally in t h e hot wa te r a t high pressure s epa ra t e s In to t h e gas phase a s boiling t raverses t h e f i r s t 30 degrees and 3 3 bars along t h e boiling curve. This concen t r a t e s t h e s t eepes t pH increase in t h e s a m e interval, result ing in a maximum r a t e of sulfide precipitation over a re la t ive ly narrow tempera ture-depth interval (Fig. 11.3b). Additional sulfide continues t o prec ip i ta te throughout boiling t o 1 0 0 ~ ~ and 1 bar, but t h e quant i t ies a r e qui te small.

Concommitant escape of H2S gas during boiling, combined with precipitation of sulfide minerals, removes aqueous sulfide, tending t o des t roy t h e gold- bisulfide complex. In a sulfide-deficient sys tem,

aqueous sulfide may be sufficiently deple ted by H2S e scape and sulfide prec ip i ta t ion t o cause gold t o prec ip i ta te along with sulfides. The requisite sulfide deficiency i s m o r e likely t o prevail in base-metal-rich waters , where prec ip i ta t ion of me ta l sulfides can remove a la rge proportion of aqueous sulfide. To be me ta l rich, t h e wa te r s must be chloride rich. Thus, deposits of gold with vein sulfides a r e likely t o have formed f rom sa l t ie r waters t han deposits of gold in near-surface hot-spring environments.

The t empera tu re dec rease t h a t accompanies boiling drives precipitation of feldspar, white mica, and chlor i te gangue a s t h e complexes of A1 and F e a r e destabil ized. Throughout a l l but t he highest t empera tu re pa r t of t h e boiling zone, t h e high pH favors format ion of gangue feldspar instead of mica. Iron-rich chlor i te fo rms ins tead of Mg-rich chlor i te because cooling destabil izes iron complexes (particularly Fe(OH)<) more sharply than Mg complexes. Also, a t t h e high t empera tu re s of t h e source fluids, t h e concentra t ion of iron is similar in magni tude t o those of magnesium (e.g., Arnorsson e t al., 1983), in t h e presence of a l t e r a t ion chlorite.

The dep th of boiling shi f t s dramat ica l ly depending on pressure. Thus, a s pressure f luc tuates owing t o self-sealing and re-breaking, t h e depth in terval of boiling changes, result ing in spat ia l overlapping of radically d i f ferent chemical regimes. The chemical d i f ferences between high- and low-pressure boiling regimes a r e exagge ra t ed by t h e d rama t i c d i f ferences between super-isoenthalpic and sub-isoenthalpic boiling (Fig. 11.7). These e f f ec t s undoubtedly contr ibute significantly t o t h e "telescoping" of zoning in epi thermal deposits and t o a l t e rna t ing o r e and was t e bands in veins. The s a m e e f f e c t s could also produce "mineralizing" and "barren" s t ages in o r e formation, a s indica ted by crosscut t ing vein relationships.

Condensation of t h e boiled gases a t 95OC produces a carbonic ac id solution with a pH of 4.5. Oxidation of t h e condensate plus remaining H2S- bearing gas by a tmospher ic O2 produces acid-sulfate wa te r of very low pH (pH 2 t o 4, depending on how much H2S oxidizes), capable of a l te r ing wall rock t o kaolinite. Oxidation of t h e boiled gas by mixing and react ion with oxygenated ground water produces a sulfate-carbonate water of near-neutral pH. The l a t t e r calculation showed t h a t acid-sulfate wa te r s cannot form by oxidation of gases by dissolved 0 because concurrent dilution precludes t h e hig{ concentra t ions of sulfuric ac id needed t o achieve very low pH. Thus, acid-sulfate wa te r s must form above t h e wa te r table , a s sugges ted by White (1957).

Mixing of acid-sulfate wa te r (or acid-condensate water ) with t h e gold-bearing liquid remaining f rom boiling causes gold t o prec ip i ta te , a s acidification drives off H2S gas and shi f t s aqueous bisulfide t o H2S, destroying t h e Au(HS)- complex. O the r meta ls t a t 5+ a r e com lexed in bisuhides, part icularly H ~ ~ ' , As , and Sbg+, also prec ip i ta te upon acidification, account ing fo r t he i r common association with gold in ac id hot-spring environments. Gold a lso prec ip i ta tes upon oxidation of Au(HS)j in t h e boiled water by dissolved O2 in ground water.

Although we have inves t iga ted in some deta i l many of t h e processes t h a t may form o re in volcanic

epi thermal environments, t h e r e a r e many processes t o examine. Among t h e m o s t pressing of t hese are: (a)boi l ing of salty, sulfide-deficient wa te r s t o inves t iga te precipitation of assemblages of sulfide minera ls with gold; (b) mixing of cold, me teo r i c ground wa te r with ho t ( 250°c), metal-bearing water without allowing boiling, t o inves t iga te th is mechanism of vein-sulfide format ion ( resul ts shown in F i . 11.14 ind ica t e t h a t i t may be q u i t e effective); and (c? boiling of metal-bearing wa te r s wi th provision f o r handling Hg, Sb, and As in t h e gas phase, t o eva lua t e t h e re la t ive impor t ance of gas fi liquid t ranspor t of t hese meta ls ; a n d (d) reac t ion of t h e various liquids and condensates with volcanic hos t rocks, t o eva lua t e more fully t h e relationships be tween wall-rock a l tera t ion and o r e fluid chemistry.

ACKNOWLEDGMENTS

We would like t o thank Bill Gemuts and t h e Anaconda Minerals Company fo r supporting much of t h e work presented here. The University of Oregon provided comput ing funds f o r t h e projec t f o r which we a r e mos t appreciative.

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proposed program can be evaluated. In practice, t he re a r e much i tera t ion and reconsideration among the s t r a t eg ic factors. Many organizations develop and exploration s t ra tegy in parallel with o ther strategies; fo r example, a s t r a t egy t o acquire identified deposits.

Organization Objectives

Th'e goal of most organizations employing

NEW DEPOSIT TYPES exploration methods is profit , which may be achieved through a variety of objectives. Examples of organization objectives t h a t employ exploration a r e l isted in Table 12.1. These objectives a r e used in various combinations t h a t r e f l ec t t h e size, maturity, and history or cul ture of a n organization. Non-profit organizations, such a s governments, employ exploration methods for land-use planning and t o Time identify sources for mineral supply.

d n exploration s t r a t egy is most likely t o achieve Figure 12"' Of i t s objectives if t h e individual objectives a r e clearly

changes in defined and communicated to the exploration with time. organization. If t h e object ives a r e unclear or

unrealistic, i t will be difficult for t h e best of exploration organizations t o be successful. For

a n assemblage of geologic d a t a or geologic d a t a and concepts. Few geologists share precisely the s a m e definit ions of model or t h e various types of models used in mineral exploration. Some geologists consider models too risky and unreliable for use in exploration. In t h e writer's opinion, problems associated with models arise from thei r imprecise definition, development, and use, not f rom t h e models themselves. The use of models in exploration, therefore , must emphasize explicit definitions and techniques t h a t promote more reliable development and application of models. The data-process-criteria model presented in the final sect ion of th is paper addresses these requirements.

STRATEGIC FACTORS

Of the numerous f ac to r s t h a t a f f e c t exploration, I 1 f ac to r s have been se lected a s most significant in t h e development of an exploration s t ra tegy. Each of these I 1 f ac to r s is shown in Figure 12.2 and briefly discussed below. Geologic information is one of t h e s t r a t eg ic fac tors , and like t h e others, i t impacts and is impacted by every other s t r a t eg ic factor. Achievement of t h e goals and objectives of an organization requires t h e balanced consideration of e a c h s t r a t eg ic factor.

Development of an exploration s t r a t egy considers t h e s t r a t eg ic f ac to r s in t h e approximate sequence presented in Figure 12.2. The objectives of t h e organization a r e f i rs t identified and potentially profitable commodities selected. The collection and in terpre ta t ion of geologic information is then coordinated with the organization's f inancial resources, staff capabilities, considerations of regulations and land availability, compet i tor ac t iv i ty , and previous exploration. Once exploration methods have been identified for a particular deposit t ype in a particular region, t he opportunities and risks of t h e

example, during t h e l a s t decade, t h e mineral exploration divisions of more than a dozen oil companies explored fo r epi thermal precious-metal deposits in t h e United Sta tes . Even though some of t h e companies were "successful" in acquiring, discovering, and even developing gold and silver properties, over half a r e now ou t of t h e mineral business, and most of t h e remaining oil companies have severely c u t back thei r exploration programs. The objectives t h a t t hese organizations pursued proved t o be unrealist ic for thei r corporations in t e rms of scale of operation, re turn on investment , and discovery potential. The reader i s r e fe r r ed t o Snow and Mackenzie (19811, Bailly (19791, and Kyle (1984) for fur ther discussions of exploration organization objectives. The exploration geologist is responsible for understanding t h e organization objectives, determining within his a r e a of exper t i se if they can be achieved, and contributing to development of t he best

DEVELOPMENT OF EXPLORATION STRATEGY,

Figure 12.2. Approximate sequence for consideration of geologic information and the other ten strategic factors in developing an exploration strategy.

Table 12.1--Examples of organization objectives that employ exploration

For Profit Organizations

Maintain organization's business

Develop feed for existing milling facility

Develop new production facilities

Increase organization's business

Develop additional production facilities

Expansion into new countries and exploration areas

Integrate organization's refining, fabricating, and marketing business back into production

Identify and acquire lands and properties with high mineral potential for sale or joint venture

Exploit existing organization assets

Forest lands, railroad right-of-ways, ranch lands, mineral holdings

Technology--exploration methods of hardware, research facilities

Information--data files for exploration regions, properties, and mineral-deposit types

Personnel--knowledgeable geologists, engineers, metallurgists, landmen, managers, negotiators, and businessmen

Nonprofit Organizations

Evaluate mineral potential of lands for land-use planning and decisions

Identify source of mineral supply

exploration strategy. Only in th is way can t h e co r rec t geologic information fo r t h e c o r r e c t deposit type(s) and t h e co r rec t exploration a rea be selected.

Commodity Pr ices

Metals prices a t t h e t i m e of production will de termine t h e profitabil i ty of most mining operations. Metal prices r e f l ec t both supply and demand, and although geologists and geologic information contr ibute l i t t l e t o predicting the l a t t e r , they do contr ibute to predicting t h e former. For example, d a t a for world gold production (Table 12.2) and geologic information for t h e gold deposits of Russia and South Africa suggest t ha t , barring polit ical changes, deposits in these countr ies will continue t o dominate world gold production and prices for t h e

foreseeable future. New gold deposits, therefore , must be of geologic types t h a t a r e profitable a t cu r ren t gold pr ices and capable of surviving periods of lower prices. This will include cer ta in bonanza and low-cost, open-pit, disseminated epi thermal deposits. The grade, geometry , tonnage, e tc . of many epi thermal deposits, however, preclude economic production a t these gold prices. Gold-price forecas t ing would b e vastly more speculative and risky, were i t not fo r worldwide geologic and economic information fo r t h e various types of gold deposits. Geologists contr ibute t o more accura t e price forecas t ing by assuring t h a t geologic information for gold deposits, deposit types, and t h e potential fo r new discoveries is a s comple te and accura t e a s possible.

The risks associa ted with commodity price forecas ts a r e considered in the selection of

T a b l e 12.2--World mine p r o d u c t i o n of gold i n 1984

C o u n t r y

South A f r i c a

USSR

Canada

Uni t ed S t a t e s

China

B r a z i l

A u s t r a l i a

Papua, New Guinea

P h i l i p p i n e s

Colombia

Othe r s

World T o t a l

P roduc t ion Thousands of Pe rcen tage Troy Ounces of T o t a l

21,905 47.9

8 ,650 18.9

2 ,614 5.7

1 ,902 4.2

1 ,900 4.1

1,611 3.5

1,199 2.6

835 1.8

773 1.7

735 1.6

3 606 I 8.0 - 45,730 100.0

Source: American Bureau of Metal S t a t i s t i c s (1985)

organization objectives and s t ra tegies for thei r achievement. The decisions t h a t result will de termine the deposit types, geologic domains, and geographic a reas appropr ia te for exploration and for which geologic information will be required.

Financial Resources

Geologic information and financing for exploration programs a r e closely interrelated. Geologic information for deposit types, proposed exploration a r e a s and exploration methods a r e used t o e s t ima te budget requirements for exploration programs. The annual funding and duration of an exploration program will, in turn, determine t h e s ize of t h e exploration organization, t h e types of t a rge t s pursued, and t h e geologic information and exploration methods used. If a n organization's resources a r e l imited, i t may require a particularly advantageous geologic opportunity (see below) to complete successfully. Finally, geologic information fo r t h e deposit t ype sought will indicate t h e probable size, complexity, and cost of development, which must be reconciled with t h e investment and risk t h e company is prepared t o make in property development. For example, in r e m o t e areas such a s portions of British Columbia, development costs may exceed a company's guidelines or capabilities. This has led one

organization (Shillabeer, 1985) t o devise a screening procedure t o identify geologically favorable a reas t h a t m e e t development cost , environmental, and land availabil i ty cri teria. Funding levels for exploration organizations may re f l ec t a mul t i tude of fac tors , including a percentage of corporate earnings, excess available funds, exploration management requests, cu r ren t me ta l markets , and t h e increasing or decreas ing confidence in and performance of t h e exploration organization.

Studies of exploration and discovery cos ts have recent ly been made for Canada (Cranstone, 1980, 1983; Mackenzie, 19841, Australia (Mackenzie and Woodall, 1983; Mackenzie and Bilodeau, 19841, t h e United S ta t e s (Rose and Eggert , 1983; Eggert , 1983), and South Afr ica (Kyle, 1984). In general, t hese s tudies indicate t h e high average cost of discovering a mineral deposit and the uneven discovery success among exploration organizations. In Australia, f o r example, d a t a for t h e period 1955-1978 showed t h a t t h e ave rage cost of finding an economic mineral deposit was $38 million, and t h a t t h e ave rage financial re turn on mineral exploration was most una t t r ac t ive (Woodall, 1984a). In order fo r mineral exploration t o be a n a t t r ac t ive business venture, therefore , a n organization must expec t and must be ab le t o achieve b e t t e r t han average results. Geologic information fo r mineral deposit types and exploration a reas is cr i t ica l

fo r es t imat ing t h e likelihood of acceptable exploration resul ts and t h e financial requirements t o achieve them.

Exploration Organization

The size, experience, and "culture" of t h e exploration organization will determine, in large par t , how i t responds t o t h e organization objectives. Some organizations seem mos t successful a t grass-roots exploration, while o the r s excel a t t he acquisition and development of properties. Western Mining Corporation, for example , is a discovery company t h a t was ins t rumental in delineating, among others, t h e Kambalda nickel deposit , t he Yeelirrie uranium deposit, and the Olympic Dam copper-uranium-gold deposit (Woodall, 1984b). Ranchers Exploration and Development Company, now wholly owned by Hecla Mining Company, has been successful in exploration but has excelled in t h e acquisition, development, and mining of properties such a s t h e Escalante epi thermal silver deposit in southwestern Utah. Some organizations a r e b e t t e r adapted to working on particular deposit types or, in particular, geologic terrains. For example, L a c Minerals has been remarkably successful in the discovery of Archean s t r a t i fo rm greenstone-belt gold deposits. According t o Valliant (19851, a familiari ty with the nature of gold occurrences in t h e Malartic area , and geologic concepts developed by Ridler (1970) and Hutchinson (1976) led successively t o discoveries a t Bousquet (6.6 million tons a t 0.14 ounces per ton Au), Doyon (zone 2 contains 8.8 million tons a t 0.149 ounces per ton Au), and Hemlo (47 million tons a t 0.174 ounces per ton Au). Distinctive capabili t ies, styles, and objectives such a s these companies display may override some of t h e o ther s t r a t eg ic f ac to r s and strongly influence t h e exploration s t r a t egy and the geologic information t h a t d i f ferent organizations use in exploration. The emphasis a t Homestake Mining Company, for example, is on maintaining and strengthening i t s position a s t h e premiere U.S. gold mining company (Anderson, 1982). Despite i t s presumed exploration advantage through ownership of t he Homestake mine in South Dakota, t he company is looking for various deposit types t h a t m e e t s i t s tonnage, grade, and earnings requirements. This program has led t o t h e discovery of t h e McLaughlin deposit in California.

Regulations and Land Availability

S ta t e , local, and federa l governments establish tax, environmental, and other regulations t h a t a f f e c t t h e availabil i ty of land for exploration and mining and t h e profitabil i ty of mining operations. Geologic information, on t h e other hand, is essential for establishing t h e mineral potential of both public and pr ivate lands for purposes of land-use planning and t h e development of regulations. Geologic information fo r t h e Thunder Mountain epi thermal d is t r ic t in centra l Idaho, fo r example, was instrumental in t h e exclusion of t h e d is t r ic t from the River of No Return Wilderness area. Geologic information for t he proposed Cabinet Mountains Wilderness a rea in Montana has demonstra ted high mineral potential fo r portions of

t h e a r e a (Banister et al., 1981). The mineral potent ia l i s now being weighed against o ther potential land uses t o de te rmine if t h e lands will be included in t h e national Wilderness system. Difficulty in bringing the significant Crandon massive sulfide deposit, discovered by the Exxon Corporation in northern Wisconsin, in to production i l lustrates t h e importance of t a x and environmental regulations in s t a t e s without strQng mining traditions. Prior t o init iating exploration in such s ta tes , geologic information and the potential fo r discovery must be carefully weighed against t h e risk of prohibitive regulations.

Compe t i to r Activity

Exploration is always compet ing with o the r organizations, and with t h e e a r t h a s well. What compet i tors do may a f f e c t t he success of one's programs by pre-empting favorable lands, hiring t h e most competent geologists, and seizing promising joint-venture opportunities. Exploration s t r a t egy should include a rea l is t ic assessment of what t he competit ion is doing with i t s geologic information; hence, what geologic information and exploration methods will be needed t o m e e t and surpass competitors. Newmont Mining Company, fo r example, discovered t h e Carlin gold deposit in 1962, and subsequently discovered Maggie Creek, Gold Quarry, and Rain. Sixteen significant sediment-hosted epi thermal gold deposits now have been discovered, and t h e Newmont deposits may contain a s much a s half t h e to t a l gold resources in t h e sediment-hosted deposit type. Any company considering exploration fo r th is deposit type will do well t o consider t h e geologic information Newmont possesses and t h e exploration use to which i t has been and is being put. To underes t imate a successful compet i tor i s a serious er ror , but any organization can be outdistanced through be t t e r use of geologic information, o ther conditions being equal.

Previous Exploration

A real is t ic assessment of t he na tu re and e x t e n t of previous exploration on proposed exploration lands will indicate what geologic information and capabili t ies (i.e., budgets, exploration methods, t i m e f rames, etc.) will be required in a new program t o improve on previous exploration and offer a reasonable chance of success. For example, in t h e ear ly 19601s, Kennecot t conducted a regional stream-sediment geochemical program within the a r e a of t h e Belt Supergroup in northern Montana and Idaho. This program is credi ted with the discovery of t h e Troy deposit (Clark, 1971) and other s t ra t i form copper- silver deposits. Recently, new concepts f o r t h e format ion of these deposits have been published (Lange and Sherry, 1983). Any new exploration programs fo r s t ra t i form copper-silver deposits in this region will have t o improve upon these ear l ier programs and concepts and whatever exploration use has been made of them. Although i t may be difficult t o obtain reliable information on previous exploration programs, such information is worth t h e considerable e f f o r t i t takes t o obtain it.

Geologic Information

Geologic d a t a and concep t s make up the geologic information of exploration. This information includes d a t a and concepts for t h e basic sciences, mineral- deposit types, geologic t e r r anes in which t h e deposits occur, and regions considered for exploration. Information f rom previous exploration programs is equally important , a s discussed in the previous paragraph. Geologic information includes all information re levant t o exploration, such a s geophysics, rock and s t ream-sediment geochemistry, topography, r emote sensing, geomorphology, and t h e concepts of processes involved in mineral deposit format ion ( to mention only a f ew examples). The information ranges in sca l e and type from mountain ranges and concepts of p la te tec tonics t o fluid inclusions and concepts of boiling and me ta l precipitation. All of i t i s potentially re levant t o exploration; much of i t ult imately will prove unnecessary, and much may in pract ice be missing. The exploration geologist uses available geologic information t o mos t ef f ic ient ly and effect ively pursue t h e organization object ives and t o identify additional information t h a t is wor th collecting.

The creat ive use of geologic information in exploration is i l lustrated by t h e discovery of the Olympic Dam deposit in South Australia by Western Mining Corporation (Haynes, 1979; Woodall, 1984b; and Lalor, 1984). In 1972, t h e company commenced a study of t h e Proterozoic and Lower Paleozoic rocks of South Australia in search for a sediment-hosted copper deposit, similar t o those in t h e Zambian Copperbelt. During t h e preceding t h r e e years, a conceptual model fo r this deposit type had been developed. The model required assembling geologic information for particular geologic character is t ics f rom t h e l i tera ture , regional reconnaissance, records of previous exploration programs fi led with t h e S t a t e Survey, regional gravity and magnet ic surveys, and l ineament studies. By July, 1975, t h e f i r s t drill hole had in tersected a blind mineralized in terval (114 f e e t a t 1.05% Cu), a t an approxoimate depth of 900 feet . Subsequent drilling led t o t h e delineation of reserves, repor ted in 1983 t o be 500 million tons of o r e grading 2.5% copper, 0.08% U308, 0.017 ounces per ton gold, and 0.17 ounces per ton silver. Olympic Dam, therefore , is a major discovery, and one of t h e world's g rea t mineral deposits.

Many geologists have tended t o a t t r ibu te t h e discovery of Olympic Dam t o luck. The discovery certainly included some unexpected, o r lucky, aspects, including t h e presence of t h e deposit in significantly older sediments than had been expected, and geologic character is t ics of t h e deposit t h a t d i f fer significantly from typical sediment-hosted copper deposits (i.e., t h e abundance of breccia and t h e high concentrations of iron, uranium, gold, and rare-ear th elements). Also, despi te t h e success of t h e f i rs t drill hole, t h e high- grade zone (2.2% Cu) was not encountered until drill hole 10 in November, 1976.

On t h e other hand, t h e f i rs t dril l hole was s i ted a f t e r t h e careful development of t h e model and in terpre ta t ion of much regional geologic da ta , and i t did discover a major deposit. The t a r g e t a r e a m e t t h e

principal requirements of t h e model including regionally a l t e red basalts, numerous copper anomalies, magnet ic and gravity anomalies and lineaments. Furthermore, within t h e Zambian Copperbelt a r e s t ra t i form iron deposits, uranium deposits, and structurally controlled gold deposits. Although the currently known m e t a l accumulations a r e not coextensive, Olympic Dam-like deposits may yet be discovered in t h e Copperbelt . In the writer 's opinion, f ine geologic work put Western Mining in t h e co r rec t area , so t h a t thei r perseverance and sha re of good luck were ab le t o ove rcome t h e equivalent amount of bad luck t h a t must a t t e n d eve ry exploration program.

Exploration Methods

Exploration methods a r e procedures and techniques such a s mapping, geochemical surveys, drilling and co re logging, airborne geophysical surveys, etc., through which geologic information is gathered and a model is applied t o a n exploration area. Exploration methods a r e specifically se lec ted to col lec t geologic d a t a in order t o t e s t for t h e essential cr i ter ia of a par t icular mineral deposit model in a particular exploration area . Exploration s t r a t egy is implemented through t h e exploration methods. Cos t ef fect ive , technically sound geologic, geochemical, and geophysical methods a r e essential for t h e success of any program, but t hey will not salvage a program for which t h e o the r s t r a t eg ic f ac to r s have not been properly assessed.

Opportunities

Every exploration program faces t h e challenge of recognizing an overlooked opportunity a t a known prospect o r discovering a completely new deposit. Exploration s t r a t egy should include, therefore , one or more speci f ic and compelling reasons why the program will succeed where o the r s have failed. These reasons a r e referred t o a s s t r a t e g i c opportunities, and they may include a new geologic concept, a n a rea t h a t has not previously been explored, o r a new exploration method. During t h e las t decade, t h e newly apprecia ted hot-spring and sediment-hosted (i.e., "Carlinn-type) type deposits offered s t r a t eg ic exploration opportunities for comparatively unexploited deposit types in the western United States. Without at leas t one such s t r a t eg ic opportunity a n exploration program is a p t t o fail.

Unique opportunities ar ise from t ime t o t ime in e a c h of t h e s t r a t eg ic f ac to r s , causing exploration t o be temporarily focused o r acce le ra t ed by t h a t factor. For example, if a company has excess profits, financial resources fo r exploration may suddenly increase, and t h e exploration program is temporarily accelera ted . Alternatively, a new geologic concept, such a s recognition of t h e geologic se t t i ng of t h e hot-spring epi thermal deposits may be perceived t o offer such opportunity fo r new discoveries t h a t exploration priorit ies a r e rearranged. Yet o ther programs become redirected in response t o breakthroughs in exploration hardware and methods. Incremental i m ~ r o v e m e n t s in airborne EM, fo r example , will p re sumad~y lead t o the discovery of severa l new massive sulfide deposits.

Risk - The final strategic factor is exploration risk,

which is assessed for the other strategic factors, both individually and collectively. Risk is typically reviewed intermittently throughout the development of an exploration strategy. A critical assessment of risk is made when the geologic information has been compiled, a model developed, and strategic opportunities and exploration methods identified. At this point, the discovery potential and exploration costs for specific areas can be roughly estimated, and management must determine if the risks warrant pursuing the program. If not, further research is initiated, or the project deferred. In some cases, even in the face of a high risklbenefit ratio, the decision is made to proceed in some modest way, leading to progressively deeper involvement in a dubious program.

Strategic factors are rarely, if ever, in balance. First, we have yet to define that balance and how to recognize it. Second, the skills and biases of organizations and individuals cause inevitable imbalances. Finally, responsibilities for the various strategic factors usually reside in different parts of an organization, thereby challenging one of our weakest organizational attributes, communication. One of the exploration geologist's responsibilities is to be familiar with all of the strategic factors and do what can be done to achieve a successful balance.

HUMAN FACTORS

A geologist's use of geologic information in exploration is not an entirely objective scientific and technical exercise. As information is accumulated, interpreted, and used, most geologists become personally involved to some extent. Pragmatic, empirical, logical, and rational performance becomes inevitably entwined with personal experiences, intuitions, biases, judgments, motives, feelings, and reactions to risks. This reflects the less than totally rational and predictable influences of the human brain, mind, and emotions, on people's work, and it is not all bad. However, research on the physiology and functions of the brain (Restak, 1984) and in behavioral psychology (Kahneman e t al., 1982) have shown that, under certain circumstances, human behavior is a t variance with empirical and rational evidence and one's best intentions. Most geologists can cite examples of jumping to inadequately supported conclusions, defending illogical positions, pursuing irrational programs, taking inordinate risks or refusing to take reasonable ones, refusing to acknowledge and correct errors, and making professional decisions for largely personal reasons. Research suggests (Kahneman e t al., 1982) people are neither entirely in control or aware of these tendencies and their consequences. Some of these behaviors may profoundly affect a geologist's use of geolog~c information and these behaviors are briefly mentioned below. Of particular importance to the use of geologic information in exploration are behaviors related to personal objectives, education and training, problem-

solving methods, intuition and creativity, uncertainty, and aversion to loss.

Personal Objectives

How successfully geologic information is utilized depends in part on the geologist's personal objectives and why he or she is an exploration geologist. Some geologists seem to be in exploration for a variety of reasons that may not necessarily contribute much to the organization's objectives, such as love of the outdoors or science, difficulty with the pre-med curriculum, math, or physics, the romance of exploration, or simply the need for a job. Such personal objectives are certainly not tailored to many exploration positions, which require an interest in and working knowledge of the diverse strategic factors discussed above and a respect for the inevitable organizational and managerial functions and responsibilities. Sometime early in their careers, therefore, exploration geologists must reconcile where they thought they were going professionally with where they seem to be headed, and decide if they like that career direction. If they decide to commit themselves to the evolving opportunities in exploration, most geological scientists can expect to learn a great deal about the nongeological aspects of planning and directing (strategy) and executing (tactics) exploration programs.

To survive in exploration, geological research must be closely integrated within successful exploration programs. The defunct mineral exploration research functions a t Shell Oil, Kennecott, Anaconda, and Exxon, to mention only a few, illustrate the liability of detached and inflated scientific functions. Geologists who prefer science must either embrace the exploration objectives or pursue careers in governmental agencies, teaching, or corporate research. Those who prefer more outdoor work have numerous options. Geologists who fail to make an overt decision and simply remain in exploration organizations risk becoming frustrated and burdensome to those exploration programs. The current educational and employment systems guarantee that a career in exploration geology will require some transition from the preparation and expectations of the academic environment to the reality of the exploration employment. How successfully geologists assess their personal objectives and capabilities and make this transition will influence their subsequent success in applying geologic information to the objectives of their organizations.

Education and Training

Geologists entering minerals exploration generally have been educated as scientists, not as explorationists, because that is what colleges and universities do. A rigorous scientific education does not prepare a geologist for the exploration environment in which success requires a working knowledge of the various strategic factors and applied exploration practices in addition to geological sciences. Geologists entering exploration must embrace a new set of values. Scientific training must

accep t a balanced, important , but nonetheless subordinate role. The e x t e n t t o which geologists can accommoda te t o this environment depends upon thei r personal flexibility and object ives and how effect ively t h e company approaches t h e training.

Once again, Western Mining Corporation is a valuable example. As described by Woodall (1984b), extraordinary success of t h e company has depended in par t on confidence in sc ience and scientists.

Ore-deposit models which have both empir ica l and theoret ica l support instill t h e g rea te s t confidence.

. . . Ear th sc ience r e sea rch and ore-deposit models a r e only re levant if they give us a sounder basis for t h a t confidence, make us bolder and more percept ive explorers, and help us t o be more confident in the recognition of e i the r t h e close proximity of ore o r a new o re environment. But we can follow knowledge and reason just so far , then comes the a c t of faith, t he leap beyond t h e sure path. Whether we a r e ult imately able t o t a k e t h a t s t e p is a t e s t of our u l t imate confidence in sc ience and ore-deposit models.

Western Mining has c rea t ed an exploration a tmosphere which exudes confidence in and respect for sc ience and scientists, supports their continued educat ion and creat ive work, and capi ta l izes on thei r productivity. Most o ther mining companies avoid even t h e mention of t h e words sc ience and scientists.

Problem Solving

Rational empiricism is our cultural method of choice for acquiring knowledge and solving problems. Throughout schooling, s tudents a r e rewarded for returning f ac tua l (empirical) and logical (rational) answers. Geologists a r e t ra ined to apply t h e sc ient i f ic method, which employs both observations and experience (empirical) and hypotheses (reasoning). The accomplishments of sc ience tes t i fy t o the power of t h e method. But, a s with any method, t he re a r e opportunities for misapplying rational empiricism.

Rational empiricism is most applicable when adequa te d a t a a r e available t o measure, describe, and define a system. Most geologic systems a r e incom- pletely documented, requiring the use of considerable judgment, inference, and intuition with the geologic information. Geologic sys tems also tend t o be - complicated, requiring generalization and simplification. Even under these conditions, skilled geoiogists may be able t o make in terpre ta t ions and predictions t h a t a r e useful in exploration. In o ther cases , i t is preferable t o collect additional information, if t h e cos t can be justified.

The m e r e use of an approach t h a t approximates t h e sc ient i f ic method may promote overconfidence in t h e d a t a and the conclusions, regardless of thei r quality. The re is a risk in using d a t a bases t h a t a r e incomplete or inaccurate, a s they may lead t o inappropriate hypotheses and conclusions. Poorly

thought ou t concepts and in terpre ta t ions may become confused with scientifically derived hypotheses. Geologic reasoning in exploration is inductive, employing d a t a from analog deposits t o def ine general character is t ics and hypotheses fo r a deposit type. Such reasoning is, a t best , inferential , suggestive, and permissive. I t requires continual tes t ing a s new d a t a c o m e available. Yet, a s discussed below, t h e r e is a natura l tendency t o place unwarranted confidence in one's in terpre ta t ions under uncer ta in ty and then refuse t o modify the in terpre ta t ions in t h e f a c e of new data. Rational empiricism and t h e sc ient i f ic method a r e applicable t o minerals exploration, but t h e hypotheses and predictions will tend t o be more uncertain and risky than in the basic sc iences for which the approaches were developed.

Most misapplications of rational empiricism a r e t h e f au l t of t h e user, not t h e approach. Over- dependence on rational empiricism may lead t o a loss of menta l flexibility and creat iv i ty (see below). There i s also t h e risk t h a t t he approach may be used a s a protect ion against cri t icism, ra ther than to promote understanding and t h e organization's objectives. Rigidly insisting on t h e rational empirical approach just t o sa t is fy a psychological need for secur i ty jeopardizes intuit ion (Goldberg, 1983).

Nothing in the foregoing is mean t t o impugn the use of rational empiricism in exploration. The predictive power of t h e approach is essential if exploration is t o benef i t more f rom existing data. The foregoing issues i l lustrate, however, t h a t rational empiricism is no panacea, and t h a t t h e r e a r e many opportunities for misapplications in mineral exploration.

Intuit ion and Creat iv i ty

Intuition and creat iv i ty a r e both potential a t t r ibu te s and th rea t s for exploration programs. Guided by existing empirical data , they can significantly extend what previously has been accomplished in exploration. Based upon incomplete or inaccurate data , or unfounded concepts, they waste t i m e and exploration funds. Both intuit ion and creat iv i ty have contributed t o exploration along with empiricism, rationalism, and luck. Most geologists have experienced those "flashes of genius" or inspirations t h a t have no obvious, rational origin or explanation, ye t move us toward t h e solution of a problem. Most geologists consider thei r geologic work t o be intuit ive and creat ive , a t l ea s t pa r t of t h e t ime. The use of geologic information in exploration re l ies heavily on empirical geologic data , i n t e rmi t t en t ly on in terpre ta t ions we make about those data , and less f requent ly on intuit ive or creat ive insights w e have about those data. Although infrequent, intuit ive breakthroughs t h a t a r e consistent with empir ica l d a t a may produce important exploration opportunities.

Intuition and creat iv i ty may become too independent, leading t o res is tance t o regimentation, documentation, personal appraisal systems, models, and even other people's ideas. There may be a compulsion t o do t h e job in our own special way and a fee l ing t h a t o ther geologists can be empir ica l and rational, and anyone can be lucky, so our creat iv i ty

and intuit ion a r e t h e bes t reflections of our specia l skills. A geologist's professional value and se l f -es teem a r e not measured by crea t iv i ty and intuition, but by how effec t ive ly he o r she balances t h e use of empiricism, rationalism, luck, intuition, and crea t iv i ty .

Some geologists resist any rigorous ra t ional empir ica l approach t o exploration a s a t h r e a t t o t he i r c r ea t iv i ty and intuition. In f ac t , rat ional empiricism should promote r a t h e r than encumber crea t iv i ty and intuition. The history of g r e a t discoveries i l lus t ra tes t h a t intuit ive breakthroughs corne t o those knowledgeable in t he i r fields, and t h a t luck comes most o f t e n t o those who have prepared themselves. Louis Pas t eu r put i t this way: "In t h e fields of observation, chance favors only t h e prepared mind." Cr i ck and Watson (Watson, 1969) discovered t h e double helix, no t p la te tec tonics , because biology was the i r f ield of life-long study. Within t h a t field, they were able t o be c r ea t ive and intuit ive because of , and in sp i t e of, their knowledge. Intuition and c rea t iv i ty s eem f i rs t t o absorb the bes t t h a t empir ica l observat ions and ra t ional in terpre ta t ions have t o o f f e r and then t o reorganize them in inspirational ways. In working geologic information, therefore , geologists mus t be willing t o d iges t and evaluate exis t ing f a c t s and in terpre ta t ions , allow thei r "processor" t o m a k e what i t can of them, and then ruthlessly eva lua t e t h e validity of t h e results and the i r value t o exploration. To indulge in t h e e levat ion of intuit ion over exis t ing empir ica l observations, logical and rational thought, and good for tune , is t o subs t i tu te personal objec t ives and emot ions fo r t h e organization's objectives. T o ignore o r t o fa i l t o promote crea t iv i ty and intuit ion is t o condemn exploration programs t o what has been known and what has been done without provision f o r t h e inevi table changes in exploration opportunit ies and c i rcumstances .

Uncer ta in ty

Exploration geology is uncertain due t o incomplete d a t a and more or less uncertain in ter - pre ta t ions fo r complex, ill-defined geologic systems. Explorationists a c c e p t t hese uncer ta in t ies a s risks in making exploration decisions. We like t o think t h a t we do t h e bes t we c a n with t h e information we have , but how good a r e our decisions? Are we really doing a s well a s can be expected , o r a r e t he re biases and e r ro r s t h a t make exploration less successful t han i t might be? Unfortunately, we lack adequa te d a t a f rom successful and unsuccessful exploration projec ts t o answer th is question. The bes t we c a n do a t t h e present t i m e is t o review some evidence of how people behave under uncer ta in ty and compare th is with our own exper ience with exploration programs.

Behavioral studies indica te t he re is something deep within t h e human mind t h a t abhors uncertainty. Kahneman il lustrated t h e point th is way (McKean, 1985):

There 's a strong overconfidence ef fec t . Suppose I t a k e you in to a darkened room and show you a c i rc le with no d is tance cues and a sk you how big i t is. You don't know whether it 's a smal l c i rc le very c lose o r a

l a rge c i rc le f a r away. If t h e mind worked like a computer , i t would say i t doesn't know t h e answer. But people always have a f i rm feeling about t h e c i rc le size, e v e n when they know they a r e probably wrong.

What you s e e he re is a classic example of how t h e human rnind suppresses uncertainty. We a r e not only convinced t h a t we know more about our politics, our businesses, and our spouses t han we rea l ly do, but a lso t h a t what we don't know must be unimportant.

Exper iments conducted by Gazzaniga and repor ted by Res t ak (1984) demons t r a t e t h e aversion t o uncer ta in ty in another way. I t has long been known t h a t t h e l e f t side of t h e brain contro ls funct ions of t h e right side of t h e body and v i ce versa. Therefore , t h e l e f t hemisphere receives informat ion f rom t h e r ight e y e and ope ra t e s t h e r ight hand. I t is also known t h a t severe epi lept ic f i t s can be a r r e s t ed by sever ing t h e corpus callosum, a bundle of nerve f ibers joining t h e two hemispheres of t h e brain. In one exper iment , a pa t ient whose corpus callosum had been severed,

. . . was shown two pictures, one projec ted t o t h e l e f t hemisphere (a chicken c law) and o n e t o t h e r ight hemisphere (a snow scene). H e was asked t o indica te f rom a ser ies of p ic tures what p ic ture h e had just seen. (He) s e l ec t ed a p ic ture of a chicken with his r ight hand and a p ic ture of a snow shovel with his left . When asked why h e had made these par t icular selections, (he) remarked: "That's easy. The chicken c l aw goes with t h e chicken, and you need a shovel t o c lean o u t t h e chicken's shed."

In th is instance, (his) l e f t hemisphere employed i t s language superiori ty t o cons t ruc t a plausible, "logical" explanation fo r t h e choices he had selected. But t h e explanation was wrong, a n e r ro r t h a t sugges ts t o Michael Gazzaniga t h a t our speech and language sys t ems routinely a t t e m p t t o in terpre t r a the r t han simply r epo r t our activit ies. If (his) explanation i s a t a l l typical of t h e r e s t of us, i t sugges ts t h a t t h e reasons we give for our own behavior may not be t h e sa l ient ones a t all.

Subsequent exper iments by Gazzaniga demonst ra ted t h a t t h e phenomena a lso occurs in normal people. The logic-language dominant l e f t hemisphere of t h e brain, t he re fo re , avoids being without a n explanation by routinely concoct ing "logical" explanations t h a t may be erroneous.

Based upon these examples, t h e i m p a c t of uncer ta in ty on a geologist's use of geologic informat ion may be considerable. Although geologists may r eve re t h e sc ient i f ic method and t h e concep t of multiple-working hypotheses, t h e d iscomfor t of uncer ta in ty promotes t h e accep tance of t h e f i rs t reasonable explanation for a s e t of data. This may

resul t in greater emphasis on explanations than on accuracy and truth. The s a m e tendency is described by Goldberg (1983):

Knowing feels good. The re is a cer ta in tension c rea t ed by ignorance, an incompleteness in a n unresolved problem . . . When t h e answer comes, t he re is a feeling of restoration.

Important insight i n t o some of the ways geologists may use geologic information under uncertainty in exploration is provided by behavioral s tudies repor ted by Tversky and Kahneman (1982). They no te t h a t people s t r ive t o achieve a coherent in terpre ta t ion of t h e even t s t h a t surround them, and t h a t cause-effect relations seem t o achieve this goal. In o ther words, people build models in which observations a r e explained a s e i ther the cause o r the e f f e c t of o ther observations. This is precisely what exploration geologists do when they in terpre t observations for a mineral deposit in t e rms of t h e processes t h a t probably fo rmed t h e character is t ics of t h e deposit. In thei r studies, Tversky and Kahneman (1982) investigated how subjects intuit ively build and use causal models, and how they respond when additional and conflicting information a r e supplied. The results a r e disquieting in t h a t they indicate people may introduce e r ro r s in the i r use of models. Their results a r e also unconformably familiar, a t l ea s t t o t h e writer.

Tversky and Kahneman (1982) found t h a t once a causal model has been constructed, t he re is a re luctance t o revise i t in t h e f ace of new data , regardless of t h e uncer ta in ty in t h e original model, and regardless of t h e veraci ty of t h e new data. Instead, t h e causal model is used t o explain t h e new d a t a with l i t t l e o r no revision of t h e model, regardless of how inconsistent and condemning t h e new d a t a may be. I t seems t h a t our highly developed explanatory skills and fluency of causal thinking inhibit t h e revision and correction of causal models in t h e f a c e of new data. This suggests t h a t exploration geologists will tend t o protect causal models from new geologic observations, using t h e model t o explain away t h e d a t a ra ther than performing the indicated revisions on t h e causal model. Somehow this s eems intuit ive and probably explains why many geologists tend t o resist t h e use of models. The work of Tversky and Kahneman (1982) also suggests corre lar ies t o th is phenomenon. First , people emphasize d a t a t h a t f i t cause-effect relations, and place l i t t l e o r no weight on other data. Second, people commonly overpredic t from highly uncertain models. Third, inferences from causes t o consequences a r e made with g rea te r confidence than inferences from consequences t o causes. All of these observations have implications for geologistst use of geologic information in exploration and encourage t h e development and use of modeling approaches t h a t suffer a s l i t t l e a s possible from t h e foregoing shortcomings. This is addressed in t h e approach t o modeling presented l a t e r in this paper.

Response t o uncertainty may re f l ec t one's rational empirical education, during which we were rewarded for having logical, f ac tua l answers t o every

question. The biggest risk t o most people is looking foolish, being humiliated, appearing stupid; we have a need t o be right. Techniques for using geologic information t h a t minimize these types of problems will improve exploration. The modeling methodology discussed in a l a t e r section has been designed t o minimize these problems.

Aversion t o Loss

People a t t e m p t t o avoid loss, even when i t can be shown t h a t a l ternat ives may ult imately lead t o g rea te r losses. "There is something in t h e human mind t h a t so abhors loss t h a t giving up some quantity of money, commodity, o r privilege is never fully offse t by an equivalent gain," according t o Kahneman and Tversky (McKean, 1985). In o ther words, loss looms larger than gain. This may explain, for example, why some exploration projects a r e kept alive by exploration managers or project geologists long a f t e r t h e accumulated geologic information should have ended them. Sponsors a r e unwilling t o accep t t h e termination and loss of a project. Legi t imate remaining opportunities need t o be pursued, but t h e human tendency t o selectively disregard information and keep ce r t a in projects going is obviously a problem.

Aversion t o loss may influence the use of geologic information in o ther ways. For example, geologists may act defensively or aggressively t o cri t icism of thei r project ou t of f e a r of losing thei r budgets, the i r projects, or t he respect of thei r colleagues. Colleagues, on t h e other hand, may overcr i t ic ize projects, ou t of f ea r t h a t they will lose standing in t h e organization if t h e projects appear t o successful. Despite our best intentions, geologic information may, f rom t i m e t o t ime, be used for personal purposes. Since these misuses a r e probably inevitable, geologic models should be designed t o minimize t h e impac t of personal motives and the o the r human factors.

DEVELOPMENT OF MINERAL-DEPOSIT MODELS

The responsibility of t h e exploration geologist is t o predic t where a mineral deposit is likely t o occur, t h e likelihood t h a t t h e deposit will be there, and t h e grade and tonnage t h e deposit is likely t o contain. In conjunction with regional geologic information, t h e mineral-deposit model is t he most powerful tool a geologist has fo r making these predictions. The predictiveness of a deposit model is derived from t h e geologic character is t ics of analogs of t he deposit type, t h e relations among these characterist ics, and the level of confidence t h e geologist has in both. A carefully const ructed model enables a geologist t o make be t t e r predictions about mineral deposits and be t t e r e s t ima tes of t h e confidence he is ent i t led t o have in those predictions. Modeling brings order out of chaos. Until geologic observations have been re la ted one t o ano the r in a causal model, they a r e just so many bits of data . In this section, modeling terminology, t he organization of geologic information for mineral- deposit types, and t h e development of mineral-deposit models a r e reviewed.

A mineral-deposit type is character ized by geologic and economic character is t ics compiled from analogs of t h a t deposit type. The character is t ics a r e sufficiently d i f ferent f rom those of o ther deposit types t o justify t h e deposit type's distinction for exploration. A mineral-deposit model is t h e representa t ion of d a t a and interpretations fo r a mineral-deposit type. Geologists have const ructed such models informally in thei r minds and more explicit ly in writ ing fo r decades. Written models soon become outdated, but they a r e necessary records fo r communication, documentation, and updating. Exploration geologists update models in thei r minds a s new information is accumulated. These t a c i t models a r e applied t o exploration a s they evolve, generally without wri t ten documentation; hence, without cr i t ica l review by the geologist o r geologic colleagues. As long a s this approach achieves the organization objectives, t he re i s n o compelling reason t o change i t . However, no t everyone builds and applies t a c i t models with equal success. Models benefit great ly f rom t h e cr i t ica l review of explicit writ ten versions by geologists and their colleagues.

The development of a mineral-deposit model is coordinated with t h e development of exploration strategy. Both benef i t from t h e development and refinement of t he other. The approximate coordination between model development s t eps and consideration of s t r a t eg ic f ac to r s i s shown in Figure 12.3. Initial selection of a deposit type for exploration is determined by (a) t h e organization objectives, (b) commodity price expectations, (c) f inancial resources of t h e organization for exploration and development, (d) skills and preferences of t he exploration organization, and (el availability of land fo r t h a t deposit type. The level t o which the model must be developed t o be successful in exploration will be determined principally by cu r ren t compet i tor ac t iv i t ies and the level of previous exploration on t h e available lands. The construction of the deposit model will be based upon existing geologic information and newly collected data , if t h e cos ts can be justified. I teration between t h e evolving deposit model and exploration methods will a t t e m p t t o assure t h a t t he methods significantly improve upon competitor 's approaches and previous exploration. Finally, evaluation of t h e model's reliabil i ty for exploration will be coordinated with consideration of opportunities and risks t h e model and t h e exploration methods present for t h e exploration strategy. The model development s teps a r e more fully discussed in t h e following section.

The mineral-deposit model is a cr i t ica l link between t h e s t r a t eg ic fac tors , in particular geologic information, exploration methods, exploration organization skills, previous exploration, and compet i tor activity. New geologic information fo r a particular deposit type, for example, impac t s t h e o ther s t r a t eg ic fac tors through thei r mutual relations t o the deposit model.

No reasonably complete o r consistent fo rma t has been developed for presenting a mineral-deposit model. Descriptions of a deposit t ype commonly ref lec t t h e author's particular expertise, or i n t e re s t in, for example, fluid inclusions, wallrock a l tera t ion, or

SEQUENCE OF MODEL DEVELOPMENT STEPS AND STRATEGIC FACTORS

Figure 12.3. Approximate coordination between consideration of strategic factors and model development steps.

regional s t ructure , ra ther than a comprehensive presentation of t h e geologic information required for exploration. Models published by t h e U.S. Geological Survey (e.g., Erickson, 1982; Cox, 19831, t he Geological Survey of Canada (Eckstrand, 19841, and t h e Geological Association of Canada (e.g., Morganti, 1981) r e f l ec t significant progress and in t e re s t in model compilations, but t hey generally a r e not yet adequa te for cu r ren t exploration. The modeling fo rma t presented in t h e final sect ion of this paper is intended t o be a s t e p toward fulfilling this need.

All models involve assumptions and interpretations. They a r e transient, ephemeral representations of what i s known a t t h e t ime, and a r e only a s reliable a s t h e d a t a and sc ience on which they a r e based. They a r e valuable springboards f o r exploration, new intuition, and improvement. However, no model i s beyond continued improvement and challenge; and, a s discussed in a n ear l ier section, we a r e f a r be t t e r a t building causal models than we a r e a t correcting, validating, and improving them.

How models a r e envisioned determines, in l a rge measure, how they a r e used and what they will contr ibute t o exploration programs. T o use a n analogy, I prefer t o think of a model not a s a destination but a s a mode of travel. I t i s not t h e wri t ten description of t h e model t h a t is so valuable; i t is a geologist's famil iar i ty and exper ience with i t s d a t a and concepts. The be t t e r one understands a model's data , in terpre ta t ions , and uncertainties, t he more predic table and reliable will be t h e model's application t o exploration. The advantage of a geologist building his or her own model is t h a t t h e model's con ten t s presumably will be well known. The danger in using someone else's model is t h a t t i m e will never be taken t o sufficiently scrut in ize i t s components t o justify t h e risk of i t s use in exploration. This, of course, need not be true. Perhaps i t is t h e inconsistent fo rma t of models t h a t discourages thei r cr i t ica l review and use. Even one's own models require persistent modification and updating, a s new d a t a and in terpre ta t ion become available.

Organization of Geologic Information

Geologic information for exploration is most conveniently organized by mineral-deposit t ype and exploration region (not considered here). Organization by deposit t ype is appropr ia te for exploration for t h r e e principal reasons. Fi rs t , valuable occurrences of metals, minerals, and rocks a r e found in reasonably distinctive geologic se t t ings , each with i t s unique assemblage of geologic character is t ics and exploration guides. Exploration for a deposit type, therefore , is tailored t o these characterist ics. Second, t h e organization's objectives commonly specify particular commodities, o r e grades, tonnages, and profitabil i t ies t ha t can be m e t only by ce r t a in deposit types. Finally, exploration fo r mineral deposits incorporates useful geologic information f rom a l l geologic disciplines, sca les of observation, and methods of collection. Thus, t h e mineral deposit itself is t he most appropr ia te f ramework fo r information organization. Within t h e deposit type, geologic information is commonly fur ther organized by subdiscipline (i.e., mineralogy, s t ructure , a l tera t ion, geochemistry, geophysics, etc.) and scale of observation (i.e., regional tec tonic sett ing, volcanic- plutonic province, vein s t ructure , fluid inclusion, etc.), which f ac i l i t a t e s application in the various s t ages and a t t h e various scales of exploration programs.

Model Terminology

Geologists differ in thei r use of t he t e rm model and in thei r definit ions of t h e various model types used in mineral-deposit work. Definitions employed in this paper ref lec t , wherever possible, t h e most common prior and current usage. For a glossary of model terms, t h e reader is referred to Adams (1985). A model represents some combination of d a t a and concepts, presented a s t ex t , formulae, graphics, and/or physical simulation. This definition recognizes t w o principal aspects of a model: i t s con ten t and i t s method of presentation. A mineral-deposit model represents d a t a and concepts of processes in terpre ted t o have formed the d a t a for a mineral deposit o r mineral-deposit type. An empirical model (synonyms include f ac tua l model and occurrence model) represents only d a t a and observations. I t does not represent in terpre ta t ions or concepts. A conceptual model (synonyms include genet ic model, in terpre t ive model, causal model, and process model) represents d a t a and interpretations. A character is t ic model represents data , with or without interpretations, fo r a particular character is t ic of a mineral deposit (examples include s t ructura l model, a l tera t ion model, and depositional model). A methodological mod& represents data , with o r without concepts, and a particular sequence of modeling s t eps t o achieve an objective. Doctors use methodological models in clinical diagnosis. Mechanics use them t o diagnose e lec t r ica l and mechanical malfunctions in automobiles. Examples in geology include exploration and evaluation models, which customarily prescribe a sequence of s t eps to achieve an objective. The data- process-criteria model (DPC model) presented in t h e final sect ion of this paper is a methodological model.

Level of Model Development

A mineral-deposit model is usually developed only t o t h e level of completeness t h a t is necessary for exploration success. In exploration, t he requirernents for a model a r e ( I ) t h a t i t improves on previous exploration on t h e proposed exploration lands, and (2) t h a t i t surpasses t h e exploration competitors. These requirements define a minimum model-development level. Sources for t he development of a model a r e the (I)ata, and (2) concepts for a mineral-deposit type, and (3) exploration methods applicable t o t h e i r field detection. The maximum development level is the highest possible model-development level supported by cu r ren t da t a , concepts, and exploration methods. The minimum and maximum levels, of course, a r e somewhat hypothetical and can be e s t ima ted only through a knowledge of t h e requirements (minimum level) and sources (maximum level) l isted above. In the preparation of a model, t h e level t o which the da ta , concepts, and exploration methods a r e developed depends upon (a) t h e level of competit ion and previous exploration (minimum development level), (b) t h e availability of data , concepts, and exploration methods fo r t h e deposit type (maximum development level), and (c) t h e purpose fo r which t h e model will be used. For example, a model prepared for mineral exploration will be developed t o a higher level than a model prepared fo r resource assessment in support of land-use planning and commodity supply studies. The cu r ren t model- development level is t he ac tua l or cu r ren t level t o which data , concepts, and exploration methods have been assembled and organized for a mineral-deposit type.

The level of development of a mineral-deposit model may range f rom a preliminary o r low-level model t o an advanced, or high-level model. A low- level model contains geologic d a t a but few, if any, concepts, and generally is based upon only a few analog deposits. As more analogs a r e identified and described, t h e d a t a base increases, permit t ing the identification of formation processes and new exploration guides in which confidence is high. This supports a higher level model. The exploration geologist mus t develop t h e model t o a level t h a t m e e t s t h e exploration requirements, but no more. The collection and in terpre ta t ion of unnecessary data , and t h e pursuit of time-consuming and unproductive sc ient i f ic studies, a r e t o be avoided. In pract ice , t he appropr ia te levels of model development is pursued by t r ia l and error. For example, exploration for sediment-hosted type deposits ("Carlin-type") has, a t various t imes, employed geologic character is t ics such a s proximity t o a thrus t faul t , lower p l a t e rocks, and jasperoids. We now know t h a t e a c h of these geologic character is t ics is informative, but t h a t successful exploration requires a higher level model with more predictive criteria.

The minimum, maximum, and current levels of model development increase with t ime, but at di f ferent rates. New data , concepts, and exploration methods may produce abrupt increases in t h e maximum development level, a s shown schemat ica l ly in Figure 12.4. Increases in t h e minimum development

LEVEL OF MODEL DEVELOPMENT

Development Level S E I e n f i f l S C o m D s f l , o n

M I N I M U M D E V E L O P M E N T L E V E L

L E V E L

Figure 12.4. Schematic representation of minimum model development level required for exploration, maximum development level supported by geologic data and concepts and exploration methods, and current model development level (Adams, 1985).

level may be caused by, fo r example, decrease in profit margin and increased competition.

The maximum development level may not be suff ic ient t o accommodate the minimum development level, if exploration has taken advantage of existing data , concepts, and exploration methods (note t w o t ime periods in Figure 12.4 where minimum level a t t e m p t s t o exceed maximum level). During these periods, exploration geologists will e i ther c r e a t e new geologic information and exploration methods through research o r postpone exploration. When t h e maximum development level exceeds the minimum development level, t h e current development level is raised above t h e l a t t e r , a f t e r which exploration should capi ta l ize on t h e cu r ren t model without fur ther inves tment in model development.

The development of a model i s discontinued a s soon a s t h e minimum development level has been reached; i.e., when t h e exploration guides a r e a suff ic ient improvement upon previous exploration in t h e exploration a r e a and upon the exploration guides being used by competitors. In pract ice , model development is generally not discontinued until readily available d a t a and concepts have been incorporated. Model development may also be discontinued, o r research expanded, when the minimum development level reaches the maximum development level; i.e., cu r ren t geologic information and exploration methods have been exploited in exploration. Discontinuation of model development for these t w o reasons is shown schematically in Figure 12.5 for t he s t eps of a data- process-criteria model (discussed in t h e nex t section). When a model ceases t o be successful in exploration (i.e., t he minimum development level rises above t h e current development level), model development i s resumed.

The concept of level of model development has been discussed above. The concept of variabil i ty between models is now introduced. Each mineral

deposit model const ructed is a unique combination of (1) t he purpose fo r which t h e model was developed, (2) t h e t h r e e sources for t h e maximum development level, and (3) t h e t w o requirements of t he minimum development level. The f ive principal f ac to r s tha t contr ibute t o variability between models a r e shown schematically in Figure 12.6 a s spokes in t h e model development wheel. The reader will recognize data , concepts, and exploration methods a s sources for t h e maximum development level, and compet i t ion and previous exploration a s t h e requirements of t he minimum development level. In Figure 12.7, model development wheels a r e shown for six hypothetical modeling situations. The spokes in t h e wheels a r e identical t o those in Figure 12.6. The re la t ive length of e a c h spoke indicates the emphasis placed on tha t f ac to r in t h e development of t h a t model. In t h e top row a r e models t h a t might have been developed by an academic research geologist, a U.S. Geological Survey resource assessment geologist, and a mineral

DPC MODEL PREPARATION AND DISCONTINUATION

~ j o p q o ~ / o ~ o ~ P r o c e s s e s

M O D E L D E V E L O P M E N T L o w L e v e l H l g h L e v e l

L E V E L

Figure 12.5. Schematic representation of steps in the preparation of a data-process-criteria (DPC) model and discontinuation of model development due to achievement of maximum or minimum model development level (Adams, 1985).

MODEL DEVELOPMENT WHEEL

L E V E L E L E M E N T S

M I N I M U M D E V E L O P M E N T I \

L E V E L R E Q U I R E M E N T S

Figure 12.6. Schematic representation of the factors that inf luence variability between mineral deposit models.

286 CHAPTER 12

VARIATIONS IN MODEL DEVELOPMENT WHEELS

E X P L O R A T I O N G E O L O G l S l

Figure 12.7. Schematic representation of variations between models developed for different purposes and exploration situations.

prospector. The research model emphasizes geologic concepts with only suff ic ient d a t a t o support t h e hypotheses and no consideration of exploration methods or f ac to r s ref lec t ing a minimum model development level. A model developed by a U.S. Geological Survey o re deposit specialist is a p t t o consider a broad d a t a base, a range of deposit format ion concepts, and exploration methods t o t h e ex ten t t h a t they (a) a r e necessary for development of t h e scientific model and (b) support resource assessment studies. The prospector's model, by contras t , places minimal emphasis on geologic concepts, but emphasizes all o ther fac tors , particularly previous exploration and competition.

The th ree hypothetical models a t t h e bottom of Figure 12.7 might have been developed by an exploration geologist fo r t h r e e distinct situations. The f i r s t wheel ref lec ts t h e l imi ted d a t a and virtual absence of concepts for a new deposit type. Since t h e deposit t ype has s o recent ly been recognized, previous exploration is not a concern, but t he re apparently a r e o ther organizations t h a t a r e exploring fo r t h e deposit type. The second wheel depic ts a model fo r a n old deposit type for which the re is considerable available d a t a and format ion concepts. The model i s t o be applied in an unexplored area; hence, previous exploration is of no concern, but t he re is no exploration competition. In t h e final wheel is depic ted t h e common exploration situation where t h e model for exploration for a well-known deposit t ype is t o be applied in a previously explored area , with considerable cu r ren t competit ion. In this l a t t e r case, s t r a t eg ic opportunities in geologic information and exploration methods a r e a lmost essential.

The development of a mineral-deposit model fo r sc ience is a special si tuation, and a researcher presumably will push t h e model t o t h e highest development level t h a t t ime, ta lent , and resources

permit. This will, from t i m e t o t ime, present exploration geologists with a sc ient i f ic windfall t h a t permits a significant increase in the current development level of a model for exploration.

DATA-PROCESS-CRITERIA MODEL

A data-process-criteria (DPC) model for a mineral-deposit t ype contains ( I ) t he geologic character is t ics of analogs of t h e deposit t ype (data), (2) t h e processes t h a t a r e in terpre ted t o have formed t h e geologic character is t ics of t h e deposit (process), and (3) t h e most reliable and informat ive geologic character is t ics (cri teria) fo r exploration. Exploration geologists have, more o r less informally, used this general approach in exploration fo r years. A DPC model d i f fers from t h e historic modeling approach in being more explicit (i.e., a consistent, wr i t ten fo rma t ) in a l l t h e modeling steps, and in i t s more rigorous selection and assignment of importance t o t h e cr i ter ia for exploration. The more explicit and rigorous DPC modeling approach is justified by (a) increasing complexity of geologic information fo r mineral deposits, (b) increasing cost and difficulty of exploration, (c) t h e need for more predictive and productive exploration methods, and (d) t h e tendency fo r human fac to r s t o erode t h e reliability and effect iveness of informally prepared models.

Most mineral deposits were formed by sequences of geologic processes t h a t generally included some combination of chemical, physical, and hydrologic processes. Each geologic process t h a t i s essential for t he format ion of a mineral-deposit type and i t s geologic and economic character is t ics is referred to, herein, a s a format ion process. Most format ion processes produce geologic characterist ics. The object ive of a DPC model is t o identify (a) a l l format ion processes for a mineral-deposit type, and (b) geologic character is t ics t h a t can be used t o determine whether o r not e a c h of t h e processes opera ted in a n exploration area. The presence o r absence of geologic evidence fo r e a c h format ion process indicates t h e favorabili ty of t he exploration a r e a fo r t h e occurrence of t h e deposit type.

The s t eps in t h e preparation of a DPC model a r e i l lus t ra ted below with simplified examples for t h e hot- spring-type epi thermal precious-metal deposit. There i s no a t t e m p t t o present a complete or even state-of- the-art hot-spring model, but only t o use this deposit t ype t o i l lus t ra te t h e preparation of a DPC model. The simplified examples a r e based upon publications fo r t h e analog deposits (see Table 12.3) and ar t ic les t h a t summarize t h e geology of this deposit t ype (Berger and Eimon, 1983; Giles and Nelson, 1982; Nelson and Giles, 1985). Only those portions of t he model useful fo r i l lustration a r e included. The DPC model preparation s t e p s (Fig. 12.8) include definition of a mineral-deposit type, compilation of analog deposits, selection of geologic data , data-process linking, identification of format ion processes, evaluation of data-process links, selection of diagnostic cr i ter ia , and evaluation of t h e model.

Table 12.3--Some ana log d e p o s i t s f o r t h e h o t - s p r i n g e p i t h e r m a l prec ious-meta l d e p o s i t type*

T o t a l Reserves Deposi t Tons Au Ag Name L o c a t i o n ( m i l l i o n s ) ( o z / t o n ) ( o z / t o n )

B o r e a l i s Nevada 2.5 0.08 0.62 S t r a c h a n e t a l . (1982) Reid (1984)

Buckhorn Nevada 5.1 0.045 0.6 Monroe (1984)

Hasbrouck Nevada 5 0.10 Graney (1984)

Round Mtn. Nevada 195 0.043 0.08 T i n g l e y and Berger (1985)

S l e e p e r Nevada 3.7 0.13 0.80 Eng inee r ing and Mining J o u r n a l (1985)

McLaughlin C a l i f o r n i a 20 0.16

Wau Papua, D i s t r i c t 0.58 S i l l i t o e e t a l . (1984) New Guinea P roduc t ion : m i l l i o n ounces Au

Iwato Japan 0.13 0.20 S a i t o and S a t o (1978)

*Analog d e p o s i t s d i s p l a y ev idence of hav ing formed w i t h i n few hundred me te r s o f t h e p a l e o s u r f a c e , a c c o r d i n g t o c r i t e r i a p r e s e n t e d i n t h e t e x t .

Definition of a Mineral-Deposit Type

A mineral-deposit t ype is character ized by geologic and economic character is t ics and format ion processes t h a t have been cornpiled from analogs of t he deposit type and t h a t a r e sufficiently d i f ferent from those of o ther deposit types t o justify t h e distinction of t h e deposit t ype for exploration. This pragmat ic definition emphasizes tha t if t h e object of exploration is a particular assemblage of geologic and economic characterist ics, and not t he character is t ics of even a closely re la ted deposit type, a distinct mineral-deposit model is required. For example, if t h e exploration objective is simply a precious-metal-bearing deposit, one deposit t ype t h a t addresses a l l epi thermal precious-metal deposits will qualify. If, on t h e o the r hand, t he exploration object ive is a very large, highly profitable underground mining operation, only models for cer ta in bonanza-vein type deposits will qualify.

The general geologic and economic character is t ics of a mineral-deposit type should be roughly defined before model development begins. The definition is accompanied by a brief summary of the geologic character is t ics t h a t distinguish i t f rom, and t h a t i t shares with, r e l a t ed o r similar deposit types. For example, hot-spring-type deposits a r e d i f ferent ia ted f rom other epi thermal deposits by a variable assemblage of geologic character is t ics t h a t suggest deposit formation occurred within a f ew hundred me te r s of a paleosurface. These character is t ics include ( I ) siliceous s in ter ,

(2) fumarol ic mineral precipitates, (3) hydrothermal eruption breccia t h a t includes vent breccia and some combination of e j e c t a blankets, fall-back breccia, and hydrofracture breccia, (4) a f r ac tu re stockwork, commonly peripheral t o the vent breccia, and ( 5 ) wallrock a l t e ra t ion cha rac te r i s t i c of ac id leaching above a ground-water table. None of these character is t ics is unique t o the hot-spring-type deposits; e a c h occurs in a t l ea s t some examples of o ther epi thermal deposit types. Also, most of these character is t ics a r e absent frorn some hot-spring-type deposits; only hydrothermal vent breccia and stockworks have been persistently identified in hot- spring-type deposits. Their absence from some deposits may be due t o erosion. Hot-spring-type deposits also di f fer from most o ther epi thermal deposits in thei r generally higher concentrations of Hg, TI, Sb, and Ba, in and above the gold-bearing zone. Hot-spring-type deposits in termit tent ly or persistently share additional geologic character is t ics with o ther epi thermal deposit types, including ( I ) f r ac tu re control of mineralization, (2) banded open-space vein filling t h a t includes abundant fine-grained quartz, pyr i te or marcas i te , and adularia, and (3) wallrock a l tera t ion t h a t includes propylitization, adularia a l tera t ion, and silicification.

Compilation of Analog Deposits

Analog deposits a r e the example deposits of a mineral-deposit t ype t h a t provide geologic and

288 CHAPTER 12

STEPS IN DPC MODEL analog deposits. ~f only t h e higher grade, larger tonnage deposits a r e of in teres t , t h e unique geologic character is t ics of those deposits must be identified so t h a t only they a r e discovered with t h e model.

D E F I N E Analog deposits for t h e hot-spring-type deposit

D E P O S I T T Y P E 5 a r e compiled in Table 12.3. Only deposits f o r which information was readily available a r e included. Prepara t ion of t h e simplified hot-spring model was

ANALOG based on publications for these analog deposits, but

P D E P O S I T S particularly on t h e summary a r t i c l e by Nelson and Giles (1985).

G E O L O G I C

I D E N T I F Y F O R M A T I O N P R O C E S S E S

D I A G N O S T I C CRITERIA

L INKING

E V A L U A T E D A T A - P R O C E S S

L I N K S

MODEL

F i g u r e 12.8. S t e p s i n the p r e p a r a t i o n of a data-process-criteria (DFC) model.

economic character is t ics for construction of a mineral-deposit model. Analog deposits should r e f l ec t both the variability of t he geologic and economic character is t ics and the most likely or typical character is t ics of t h e deposit type. The more analog deposits t h a t a r e included in a model, particularly from widely separa ted areas, t he greater will be t h e model's reliabil i ty for exploration. As f ew a s a dozen analogs may suffice, if t he character is t ics display l i t t l e variability. If, on the other hand, t h e analogs show considerable variability in grade, tonnage, s t ructura l control, al teration, etc., a f ew dozen analogs may be required t o develop a reliable model. High variabili ty also may indicate t h a t t h e se l ec t ed deposit t ype should be divided into two or more sepa ra t e deposit types. All readily available analogs should be included in the model.

Both t h e geologic and economic character is t ics must be compiled from, and representa t ive of, t h e s a m e group of analog deposits. Deposits discovered in exploration can be expected to have about the s a m e grade and tonnage distribution a s occurs among t h e

Select ion of Geologic Da ta

Geologic d a t a for construction of a mineral- deposit model a r e derived f rom analog deposits and geologic environments in which t h e deposits occur. Fac tua l da t a a r e developed through mapping, thin sect ion study, chemical analyses, etc., and is initially incorporated in to t h e model without interpretation. Sources of da t a include t h e l i tera ture , professional colleagues, and visits t o t h e more accessible, exposed, and controversial deposits t o t e s t t h e assembled d a t a and concepts and col lec t missing data.

The collection of geologic d a t a from t h e analogs involves both judgment and bias. A conscious e f f o r t is made t o col lec t relevant d a t a of a l l types (i.e., s t ructure , geomorphology, petrography, etc.) and scales (i.e., regional t ec ton ic se t t i ng t o deta i led deposit character is t ics) and t o avoid t h e tendency t o overemphasize t h e collection of d a t a types t h a t a r e famil iar t o t h e observer.

D a t a collection is guided by hypotheses fo r t he format ion of the geologic character is t ics of t he deposit type. This approach permits more ef f ic ient tes t ing of hypotheses during t h e d a t a collection and minimizes the collection of unnecessary da ta , even where t h e init ial hypotheses prove t o be wrong. I tera t ion between d a t a collection, hypotheses, and t h e tes t ing of hypotheses proceeds until t h e d a t a a r e exhausted, confidence in t h e format ion processes becomes sufficiently high, or t h e model is discontinued because of loss of interest .

The geologic da ta may be organized in a var ie ty of ways. Tables or ma t r i ces in which both t h e d a t a and t h e analog deposits a r e shown have been useful. The d a t a base should r e f l ec t both t h e variability among geologic character is t ics and the most typical characterist ics.

Some s t ructura l geologic d a t a for s e l ec t ed analogs of t h e hot-spring-type deposit a r e compiled in Table 12.4. The d a t a a r e incomplete and simplified and a r e presented for i l lustrative purposes only. In a complete model, a l l geologic character is t ics of t h e analog deposits, including regional geologic sett ing, host rocks, intrusions, a l tera t ion, etc. would be presented in a similar matrix-like table.

During d a t a accumulation, geologic character is t ics may be identified t h a t appear t o offer a significant advantage in exploration. Such geologic character is t ics can be immediately applied t o exploration, if confidence is high t h a t t he potent ia l value of their use exceeds the risks of their use prior t o completion of t h e model.

Table 12.4--Selected structural geologic data for some analogs of the hot-spring type deposit

Regionall District Structure

Round Data General Description Borealis Hasbrouck Mountain - W au McLaughlin -

References Berger and Eimon (1983); Reid (1984); Graney (1984); Tingley and Sillitoe Nelson and Giles and Nelson (1982); Strachan Tooker (1985) Berger (1985); et al. (1984) Giles (1985); Nelson and Giles (1985) (1985); Tooker (1985) Tooker (1985).

Tooker (1985)

Fracture zones, faults, N-trending N-trending NW- and NE- Low-angle Not described. caldera ring and graben normal faults steeply trending, normal fault fractures, joints, and and NE- and dipping faults; steeply dip- and Maar ring areas of doming and E-trending caldera ring ping faults, fault intense fracturing. shear zones fractures ( ? ) fractures and

low-angle fractures; joints

Hydrothermal Eruption Breccia

Ejecta Blanket and fallback breccia

Vent breccia and hydrofracture breccia

Stockwork

Ejecta breccia blanket up to several meters thick, generally restricted to immediate vent area but covering up to a square mile; fragments may show episodic silicification prior to brecciation; may be eroded in fossil systems.

Pipelike, sheetlike, tabular, dikelike and podiform masses of breccia; millimeters to tens of feet in thickness and diameter; pebble dikes and brecciated vein walls; isolated to multiple overlapping or neighboring breccia bodies; variably matrix and clast supported; very fine- grained to boulders.

Vein stockworks, generally in previously silicified and/or adularia-altered wallrock.

Possibly deposited contemporaneously with talus deposits

Not described Not described Thickness up Not described to 60 m

Multiple stages Irregular Linear and Irregular Present bodies in irregular veins and silicified pipelike pods rock bodies related

to fractures and fracture intersections; multiple stages

In silicified In silicified Closely related Some within Present breccia rock and to fractures hydrothermal

silicified and breccias breccias breccias

Data-Process Linking

Once geologic d a t a for t he analogs have been se lected, t h e d a t a a r e used t o identify probable processes t h a t formed t h e deposit type. The processes a r e identified by linking e a c h geologic datum t o every possible process by which i t may have been formed. Although th is data-process linking s t e p precede t h e final selection of formation processes, i t i s more conveniently discussed in t h e following section.

hydrologic processes opera ted and in a particular sequence. Each process t h a t is essential fo r some significant geologic o r economic cha rac te r i s t i c of t he deposit t ype is referred t o a s a format ion process. The absence of even one format ion process precludes format ion of t h e deposit t ype a s i t is presently defined. The objective in modeling is t o identify t h e format ion processes; t h e objective in exploration is t o identify a r e a s where a l l of t h e format ion processes operated. Since the format ion processes generally have long since ceased t o opera te , they mus t be

Identification of Formation Processes identified by the geologic character is t ics they produced. In t h e DPC modeling process, t h e format ion

The format ion of most mineral deposits requires processes a r e identified from t h e geologic data . Then t h a t a particular sequence of chemical, physical, and the most reliable, informative, and easily measurable

geologic character is t ics (diagnostic cr i ter ia) a r e then se l ec t ed for t h e ident i f ica t ion of where, in an exploration area , t he format ion processes operated.

The identification of a format ion process may need t o be augmented by informat ion on t h e intensity, duration, and repetit ion with which the process operated. In a hot-spring-type deposit, for example, o r e grade may be determined, in par t , by t h e intensity of f rac tur ing and t h e volume of hydrothermal fluid flow. Evidence fo r (a) t h e number of episodes of brecciation and (b) t h e e x t e n t of a l tera t ion of breccia f ragments associated with e a c h episode could provide information on these t w o processes.

The identification and evaluation of formation processes probably will require, in addition t o skills a t developing and assembling field and laboratory data , a working knowledge of chemical , physical, and hydrologic processes. F e w geologists enjoy such a broad, formal education; hence, t he re is a continual need t o learn during employment and t o collaborate with knowledgeable colleagues. Geologists inexperienced with these processes may find i t helpful t o build models in t e a m s with colleagues whose skills augment thei r own o r they may re s t r i c t the i r modeling t o empirical models. The misinterpretation of format ion processes through t h e misapplication of chemical, physical, and hydrologic processes not only wastes valuable t ime, but risks substi tuting erroneous conceptual programs for more reliable, empirical ones.

Formation processes a r e identified for a l l t he geologic da ta col lec ted f rom t h e analog deposits. Each geologic observation is linked t o one o r more processes, from which i t might have formed. Linking is recorded in a diagram t h a t becornes an integral par t of t h e model. Selection of format ion processes can be i l lustrated with a simplified example for t h e hot- spring-type deposit. F igure 12.9 presents geologic d a t a fo r analogs of t h e deposit type t h a t includes a breccia and geologic character is t ics A, B, C, D, and E. Possible processes t h a t may have formed t h e breccia have been identified. These possible processes include intrusion, hydrothermal eruption, sedimentation, and tectonism, and e a c h is linked t o breccia. Character is t ics of t hese particular breccias, including (a) thei r generally circular shape in m a p view, (b) sharp ver t ica l con tac t s with enclosing rocks, (c) absence of an igneous component, and (dl multiple a l t e red and breccia ted c las ts suggest t h e breccias were formed by episodic hydrothermal eruption. Hydrothermal eruption has been se lected, therefore , a s t h e preferred format ion process for t h e breccia. Tectonism is t h e most likely a l t e rna te process, because all analog deposits a r e associated with f au l t s and f r ac tu res of various orientations. I t may subsequently be shown t h a t tectonism is also a preferred process, perhaps an an teceden t t o hydrothermal explosion. Note t h a t in this simplified example, hydrothermal explosion was also supported by geologic da ta A and C. Processes se l ec t ed a s preferred processes a r e commonly, but no t always, supported by multiple geologic data. Conversely, possible processes supported by multiple d a t a f rom t h e analog deposits a r e not always judged t o be essent ia l t o t h e format ion of significant geologic and/or economic character is t ics of t h e deposit type. The

possible processes intrusion, sedimentation, and tectonism were not se lec ted a s preferred processes for t he format ion of the breccia, but in a complete model they may have been se l ec t ed a s preferred processes for o ther geologic character is t ics of hot-spring deposits. The processes me ta l leaching, a l tera t ion, and me ta l precipitation would have been supported by numerous additional d a t a in a complete hot-spring model.

In Table 12.5, t h r e e preferred processes and their a l t e rna te processes for t h e simplified hot-spring model a r e compiled. Such a table presents t h e ra t ionale for selection of t h e preferred processes and is an important component of a data-process-criteria model.

Once t h e preferred processes have been se lected, they a r e summarized in a diagram of data-process links. St ructura l geologic data , previously compiled in Table 12.4, a r e linked t o t h r e e probable formation processes fo r t h e simplified hot-spring model in Figure 12.10. Each line f rom a geologic datum t o a process

SELECTION OF FORMATION PROCESSES

GEOLOGIC POSSIBLE PREFERRED DATUM PROCESS PROCESS

ALTERNATE PROCESS

M e t a l A leaching l eh la~ th~~g

B 4 Intrusion

Brecc la Hydro therma l & Hydro therma l Tectonlsm e x p l o ~ ~ o n explosion

S e d i m e n t a t ~ o n

Tec ton ism

A l te ra t ion A l te ra t lo"

M e t a l - M e t a l p rec lp l t a t ion p rec lp l t a t lon

Figure 12.9. Simplified example of the selec- t ion of formation processes for the hot- spr i ng-type deposit.

SOME DATA-PROCESS LINKS FOR THE HOT-SPRING MODEL

G E O L O G I C D A T U M

N o r m a l F a u l t

F r a c t u r e Z o n e

C a l d e r a Ring F r a c t u r e

M a a r R i n g F r a c t u r e

E j e c t 8 B l a n k e t 1 F a l l b a c k B r e c c i a

P R E F E R R E D F O R M A T I O N P R O C E S S

H f g h H e a t F l o w

V e n t I H y d r o f r a c t u r e B r e c c i a

Figure 12.10. Data-process links for selected s t ructura l data and three pss ib le forma- t ion processes for thehot-spring deposit d e l .

Table 12.5--Three preferred and alternate processes for the simplified hot-spring model

Preferred Alternate Process Process(es) Comments

High heat flow

None

Episodic hydrothermal eruption

Episodic convection of hydrothermal fluids

Intrusion

Tectonism

Sedimentary brecciation

Indicated by the effects of hydrothermal fluids including brecciation, alteration, mineralization, and metallization; may be due to intrusion or, less likely, high regional heat flow.

Given the geologic data, no alternate to high heat flow can be imagined.

Indicated by episodic brecciation and alteration.

A probable process elsewhere in model as source of high heat flow; unlikely as alternate process for hydrothermal eruption, because of absence of igneous components in breccias and numerous effects of hydrothermal fluids.

A probable process elsewhere in model as cause of faulting and fracturing; unlikely as alternate process for hydrothermal eruption because of shapes of breccia bodies and lack of relations to significant faults and fractures; possibly an antecedent process for hydrothermal eruption.

Unlikely because of breccia shapes and crosscutting relations with wallrocks and associated hydrothermal effects.

Indicated by episodic brecciation, alteration, mineralization, and metallization.

Ground-water Unlikely that nonhydrothermal fluids circulation produced brecciation, alteration,

mineralization and metallization because of metal solubility considerations and fluid-inclusion data.

ref lec ts t he in terpre ta t ion t h a t t h e datum formed f rom t h e process, and t h a t t h e datum is a field indication fo r t h e process. In a complete hot-spring model, t h e s t ructura l d a t a shown would also link t o some processes not shown, and d a t a not shown would link t o some of t h e processes shown. During t h e identification of t h e format ion processes, a l l geologic d a t a assembled for t h e analogs should be linked t o processes. To help avoid p rematu re and prejudicial selection of format ion processes, no d a t a should be eliminated a t this s t age becuase i t is thought t o be

re la ted t o nonessential processes. Processes t h a t a r e l a t e r found t o be nonessential t o deposit formation can b e e l iminated f rom t h e model.

The identification of format ion processes may b e aided by some additional techniques. For example, thinking through t h e sequence me ta l source-transport- precipitation-enrichment-preservation may help identify additional and less obvious processes. Another technique is t o th ink of processes a s having resul ted f rom o r having been made possible by an teceden t processes and a s having produced o r permit ted

subsequent processes, a n d then a t t e m p t t o identify these additional processes. For example, consider a n enargite-gold epi thermal ore shoot. Processes an teceden t t o t h e hydrothermal process t h a t precipi ta ted t h e me ta l s probably included some combination of regional t ec ton ic processes, intrusion of a pluton, development of a ground-water source, development of hydrothermal circulation cells, f r ac tu r ing of t h e host rock, and so forth. Processes subsequent t o me ta l precipitation may have included hydrothermal leaching and redistribution of mineralization, enr ichment of t he vein below a weather ing or leaching zone, displacement of t h e vein along faul ts , and t h e format ion of topographically e levated erosional remnants of silicified wall rock. Only an teceden t and subsequent processes t h a t a r e essent ia l t o deposit format ion a r e included in t h e final model.

Finally, a glossary of geologic processes, preferably arranged by categories, may help t h e search fo r essent ia l format ion processes. For example, t h e identification of processes t h a t may have precipi ta ted precious me ta l s in a n epi thermal deposit might b e aided by r e fe rence t o a glossary of processes t h a t included (a) changes within the hydrothermal fluid, such a s boiling, mixing, and cooling, and (b) a var ie ty of wallrock reactions.

Data-process linking promotes identification of a l l possible and then probable and a l t e rna te processes t h a t could reasonably account for e a c h geologic cha rac te r i s t i c of t h e analog deposits. When t h e linking has been completed, confidence will be highest in t h e possible processes t h a t a r e linked to several geologic character is t ics . Confidence also will be highest in processes tha t a r e chemically, physically, and hydrologically reasonable in t h e envisioned deposit- forming environment. These processes a r e se l ec t ed a s preferred processes. I t is advisable t o re ta in a t l ea s t one a l t e rna te process for e a c h probable process, and more a l t e rna te~processes for more uncertain preferred orocesses. Maintaining t h e best ~ o s s i b l e a l t e rna te

.2

processes sharpens conceptualization of a l l processes, result ing in t h e selection of preferred processes t h a t a r e mos t likely and most reliable. Without continual challenge, ill-defined or inaccurate preferred processes will persist in a model.

Geologic and economic character is t ics produced by format ion processes may be e i ther relatively numerous and obvious or f ew and obscure. A t t h e s a m e tirne, processes responsible for geologic and economic character is t ics of analog deposits may be identified with re la t ive ease or with g rea t difficulty. For example, processes t h a t control barren versus mineralized veins, high-grade zones and large- versus small-tonnage deposits may be difficult t o identify. I t is important , nonetheless, t o a t t e m p t t o identify a l l essential format ion processes and cr i ter ia by which they can be recognized, but particularly those t h a t produced t h e significant economic character is t ics of t h e deposit type.

The identification of formation processes enta i ls g rea t e r uncertainty than any other s t e p in t h e construction of a DPC model. Because t h e format ion processes s o significantly a f f e c t t h e reliabil i ty of a model fo r exploration, particular e f fo r t i s made t o

eva lua te t h e reliability of t h e format ion processes. This i s customarily done a f t e r t h e format ion processes have been se lected and again during t h e final model evaluation step. The evaluation of the format ion processes a t t e m p t s t o determine t h e e x t e n t t o which the following conditions have been met: (1) a t l ea s t one process has been identified for every geologic character is t ic compiled f rom t h e analog deposits; (2) e a c h formation process is essential t o deposit formation, a s evidenced by i t s data-process links; (3) e a c h process is confirmed, where possible, by multiple data-process links; (4) techniques such a s a glossary of process terms, sequential process linking, and antecedent-subsequent processes were used t o develop t h e l ist of preferred and a l t e rna te processes; ( 5 ) processes were i tera t ively t e s t ed agains t da t a collection; and (6) e a c h process i s sound scientifically and in t h e geologic environment in which is i t in terpre ted t o have operated.

Evaluation of Data-Process Links

Application of a mineral-deposit model in exploration requires geologic cr i ter ia t h a t reliably indicate the likelihood for t h e occurrence of t h e deposit t ype in an exploration area. In t h e case of a low-level, empirical model, t he c r i t e r i a a r e empirical geologic character is t ics t h a t occur with t h e analog deposits. In a higher level or more conceptual model, t h e c r i t e r i a a r e also geologic character is t ics of t h e analog deposits, but they have been shown t o have formed f rom particular deposit format ion processes. These geologic character is t ics a r e referred t o a s diagnostic cr i ter ia , and e a c h one provides information on t h e presence or absence of one o r more format ion processes in a n exploration area. The evaluation of geologic data /character is t ics for se lec t ion of diagnostic cr i ter ia i s discussed below. Exploration use of empirical observations fo r lower level models has been common in exploration and is no t discussed in this paper.

Once format ion processes have been identified, a s described in t h e previous section, e a c h geologic character is t ic linked t o them is evaluated fo r t h e s t r eng th of t h e evidence i t provides fo r t h e process and, therefore , suitabil i ty of t he character is t ics for a diagnostic criterion. To qualify a s a diagnostic cri terion, a geologic character is t ic mus t provide s t rong evidence t h a t a format ion process did or did not ope ra t e in a n exploration area . Evaluation of t h e data-process links is, therefore , t h e next s t e p in model development.

Some geologic character is t ics a r e formed only by o n e geologic process and a r e always found with t h a t process. Such character is t ics a r e particularly valuable evidence for t h a t process because, if present, they indicate t h a t t h e process opera ted and, if absent , they indicate t h a t t h e process did not operate. One may say, therefore , t h a t they a r e necessary evidence for t h e process t o have opera ted and suff ic ient evidence t h a t i t did. Other geologic character is t ics a r e always fo rmed by a particular process, but they a lso a r e formed by other processes. These character is t ics a r e necessary, but less sufficient, evidence for t he process. Character is t ics t h a t a r e not a lways formed

by a par t icular process and a r e not formed by o the r processes a r e suff ic ient but less necessary evidence fo r t h a t process. Finally, character is t ics t h a t a r e not a lways formed by a par t icular process and a r e formed by o the r processes a r e both less necessary a n d less suff ic ient evidence f o r t h a t part icular process. The task, therefore , is t o identify t h e most necessary and mos t suff ic ient geologic character is t ics for e a c h format ion process. These character is t ics a r e s e l ec t ed a s diagnostic c r i t e r i a because they a r e t h e mos t informat ive and re l iable cr i te r ia for de termining if a format ion process ope ra t ed in a particular area . Information on a l l t h e format ion processes will, in turn, indica te t h e likelihood t h a t a deposit will occur in a n area.

Es t imat ion of t h e necessity and sufficiency of a cr i te r ion fo r a format ion process employs t h e rules shown in Table 12.6. Necessity and sufficiency a r e evaluated semiquant i ta t ive ly (i.e., high, in termedia te , and low) ins tead of numerically, because the e s t i m a t e s a r e inexact and subjective. In Figure 12.10, s t ruc tu ra l d a t a a r e linked t o t h r e e prefer red format ion processes f o r t h e hot-spring deposit- type model. On t h e l e f t side of Figure 12.11, t h e necessity and sufficiency of e a c h of t hese links has been e s t ima ted and placed nex t t o t h e link. Each geologic da tum with high necess i ty and/or high sufficiency for one or more process has been se l ec t ed a s a diagnostic cr i te r ion and l i s ted on t h e right s ide of Figure 12.11. L e t us now review how necessity and sufficiency were e s t ima ted fo r s o m e of t hese links.

Caldera r ing f r ac tu re s a r e impor tant indica tors fo r some processes no t shown in Figure 12.10; for example , intrusion, volcanism, and fracturing, but they a r e also linked t o high h e a t flow. We will eva lua t e t h e

SOME NECESSITY AND SUFFICIENCY ESTIMATES AND DIAGNOSTIC CRITERIA FOR THE HOT-SPRING MODEL

P R E F E R R E D G E O L O G I C D A T U M F O R M A T I O N P R O C E S S O l A O N O S T I C C R i T E R l O N

H O h n e a t i t o w

E . , S O d C n y r r o l h e r r n a l D r Y P l O "

E P s o d l r h Y d r * , n e r m a i l , " , d , l o *

Figure 12.11. Necessity and sufficiency estimates for selected data-process links and selected diagnostic criteria for the hot- spring deposit model.

necessity of this link, i.e., t h e necess i ty of a ca ldera ring f r a c t u r e a s evidence t h a t high h e a t f low occurred in a n area . Using t h e ru le fo r necessity in Table 12.6, i s a ca ldera ring f r ac tu re formed by or associa ted with high h e a t flow consistently, in some/many cases o r rarely? I t was decided t h a t ca ldera ring f r ac tu re s occur in a r e a s of high h e a t flow in some lmany cases; hence, i t s necessity is in termedia te . In o the r words, in only some/many cases do a reas t h a t show evidence of high h e a t flow (defined for our purposes a s evidence of

Table 12.6--Rules for the estimation of necessity and sufficiency of a geologic datum or characteristic for Lhe occurrence of a formation process to which the datum is linked in a data-process-criteria model

If the criterion is formed by or - associated with the process:

Consistently

In some/many cases

Rarely

If the criterion (a) is produced - only by the process (i.e., not produced by plausible alterna~ive processes) and (b) reflects the significant aspects of the process:

In somelmany cases

The Necessity is:

High

Intermediate

Low

The Sufficiency is:

High

IntermediaLe

Low Rarely

shallow intrusion, hydrothermal a l te ra t ion , hydrothermal minera l iza t ion and metall ization, etc.) occur with ca ldera ring f rac tures . Therefore, a ca ldera r ing f r a c t u r e is n o t highly necessary fo r high h e a t f low t o have occurred. Now l e t us e s t i m a t e t h e sufficiency of t h e link, using t h e sufficiency rule in Table 12.6. I t was concluded t h a t a ca ldera ring f r a c t u r e i s consistently produced only in a r eas of high h e a t flow; hence, i t s suff ic iency f o r high h e a t flow is high. In o ther words, ca lde ra ring f r ac tu re s a r e found in a r e a s t h a t consistently show evidence of high hea t flow; hence, t h e presence of a ca ldera ring f r ac tu re is compelling evidence (i.e., highly suff ic ient evidence) t h a t high h e a t flow occurred in t h e area. T h e absence of a ca ldera ring f r a c t u r e is negat ive but ambiguous evidence fo r t h e absence of high h e a t flow, whereas t h e presence of a ca lde ra r ing f r ac tu re is highly suff ic ient evidence f o r t h e occu r rence of high h e a t f low in a n area .

Consider next t h e link f rom stockwork veining t o high h e a t flow. I t was concluded t h a t stockwork veining rare ly occurs in a r e a s of high h e a t flow; hence, i t s necessity fo r t h e process is low. The sufficiency of stockwork veining f o r high h e a t f low is high, however, because stockwork veining i s rare ly found in a r eas t h a t have no t exper ienced high h e a t flow. The link be tween stockwork veining and episodic hydrothermal fluid flow has i n t e rmed ia t e necessity, because stockwork veining is associa ted with th is process in only some/many cases. The sufficiency of s tockwork veining i s a lso in termedia te , because s tockwork veining is associa ted with a r eas t h a t show no evidence of episodic hydrothermal fluid flow in some/many cases. By contras t , vent /hydrofrac ture breccia is consistently associa ted with t h e hydrothermal eruption process and does not form by o the r processes; hence, i t s necessity and sufficiency a r e high. The o the r links have been evaluated in a similar manner.

Selection of Diagnostic Cr i t e r i a

Geologic d a t a in Figure 12.11 have been se l ec t ed a s diagnostic c r i t e r i a t h a t have high necessity and/or sufficiency fo r o n e o r m o r e processes. These geologic character is t ics a r e mos t informat ive fo r de termining if a format ion process ope ra t ed in a n exploration area . Note t h a t t h e e j e c t a blanketffallback and vent /hydrofrac ture breccias were no t s e l ec t ed a s c r i t e r i a fo r episodic hydrothermal fluid flow. Although these geologic cha rac t e r i s t i c s a r e highly suff ic ient fo r hydrothermal eruption, t hey do not provide d i rec t informat ion on hydrothermal fluid flow, fo r which t h e evidence must indica te la rge volumes of fluid flow. In a comple t e model f o r t h e hot-spring deposit type, diagnostic c r i t e r i a f o r hydrothermal fluid flow would include (a) e x t e n t of wallrock a l tera t ion , (b) s t ages of wallrock a l tera t ion , (c) open-space mineralization, a n d re la ted cr i te r ia . The episodic na tu re of hydrothermal eruption and hydrothermal fluid flow would a lso have t o be conf i rmed by speci f ic evidence, such a s multiple s t ages of crosscut t ing f rac tures , banded mineralization, episodically a l t e r ed wallrocks, and so forth.

The mos t useful diagnostic c r i t e r i a fo r exploration a r e those t h a t (1) have high sufficiency

and/or necessity, and (2) a r e easily and inexpensively measured by explora t ion methods. Cr i t e r i a t h a t have high necessity and sufficiency f o r a process may be a l l t h a t is required t o de t e rmine if t h e process operated. In prac t ice , i t i s bes t t o confirm t h e presence o r absence of a process wi th a t l ea s t t w o cr i te r ia . Some c r i t e r i a have high necess i ty and/or sufficiency fo r more than o n e process; hence , t hey a r e especially useful.

Evaluation of Data-Process-Criteria Model

The sys t ema t i c evaluation of a DPC model is t h e l a s t and mos t impor t an t s t e p in model development. As discussed ear l ie r under Human Factors , t he re is a na tura l human tendency t o build causal models without adequa te regard fo r the i r validity. To minimize this risk, no model should be applied t o exploration without an expl ic i t assessment and documenta t ion of i t s reliability, hence t h e risks associa ted with i t s use. The wri ter has developed a ser ies of t e s t s t h a t evaluate each s t e p in model development. S ince a l l modeling steps, including t h e se lec t ion of geologic d a t a for analogs, include subjec t ive judgment, evaluation of a model and i t s suitabil i ty fo r exploration is also subjective. The evaluation t e s t s a r e briefly summarized below.

Each of t h e seven previous modeling s t eps (Fig. 12.8) a r e evaluated fo r accuracy, completeness, and t h e conf idence t h a t s eems warranted by the results of t h e step. Tes t s f o r accuracy emphasize t h e verac i ty of (a) t h e uniqueness of t h e deposit type, (b) t h e appropr ia teness of t h e analog deposits, (c) t h e geologic and economic da t a , (dl t h e sc ient i f ic principles used in se lec t ing t h e format ion processes, (el t h e prefer red and a l t e r n a t e processes, (f) e s t ima te s of necess i ty and sufficiency, and (g) t h e diagnostic c r i t e r i a se lec ted .

Tes t s f o r completeness de t e rmine if (a) geologic and economic s imi lar i t ies among and di f ferences between t h e deposit t ype and r e l a t ed deposit types a r e a s complete ly documented a s cu r r en t d a t a permi ts and exploration warrants; (b) a l l avai lable analog deposits have been compiled; (c) a l l types (mineralogical, tec tonic , s t ruc tura l , s t ra t igraphic , etc.) and re levant sca les (regional t o microscopic) of geologic d a t a for t h e analogs a r e represented; (d) analogs accu ra t e ly r e f l ec t t h e range of geologic and economic character is t ics of t h e deposit type; (e) prefer red and a l t e r n a t e processes have been identified fo r all

cha rac t e r i s t i c s of t h e analog deposits; R0A2%sity and suff ic iency have been e s t ima ted for a l l data-process links; and (g) if t h e most informat ive (highest necess i ty and/or sufficiency) and reliable geologic d a t a have been se lec ted a s diagnostic cri teria.

Conf idence in t h e resul ts of e a c h modeling s t e p i s based upon t h e t e s t s fo r completeness and accuracy and ce r t a in addit ional f ac to r s fo r some steps. Fo r example , t h e reliabil i ty of a format ion process depends, in l a rge par t , on t h e number of d a t a t o which i t is linked, a n d genera l sc ient i f ic and geologic conf i rmat ion of t h e process, including t h e conditions under which i t opera tes , how i t opera tes , and t h e geologic cha rac t e r i s t i c s i t produces under various

geologic conditions. Confidence in t h e diagnostic cr i ter ia depends upon t h e reliabil i ty of their necessity and thei r sufficiency e s t ima tes and upon t h e reliability of t h e processes themselves.

Every mineral-deposit model enta i ls uncertain- t i e s due t o incompleteness and t h e possibility of er rors in d a t a and interpretations. A carefully prepared and evaluated data-process-criteria model permits evaluation of the uncer ta in t ies and more responsible exploration decisions. The a l t e rna te approach, namely t h e use of se lec ted empir ica l observations, also enta i ls uncertainties, but without t h e benef i t of knowing where in t h e model they are , the i r probable significance , and how they might be minimized. Existing d a t a and concepts will not permit t h e development of high-level models fo r a l l deposit types, necessitating t h a t exploration e i the r proceeds with empirical observations or awa i t s t he development of new d a t a and concepts. Even when only empirical o r low-level models have been prepared, they a r e available t o guide continued research and capi ta l ize on advancements a s they occur.

Application of the Data-Process-Criteria Model t o Exploration

Once a model has been evaluated and judged sufficiently reliable fo r use in exploration, specific exploration methods a r e se lec ted for applying the model to exploration areas (see ear l ier discussion under St ra tegic Factors). Exploration methods a r e se lec ted t o provide t h e most rapid and inexpensive field and laboratory da ta fo r each diagnostic criterion. Methods applicable to exploration for hot- spring deposits might include, for example, (a) use of existing regional mapping t o identify volcanic-plutonic provinces, s t ructures (calderas, faults, f r ac tu re zones, etc.) and a l tered zones; (b) ground reconnaissance t o identify individual hydrothermal sys tems (a l tera t ion and mineralization) and evidence of near-surface hydrothermal processes (sinter, e j e c t a blankets, fallback breccias, vent breccias, hydrothermal breccias, stockworks, etc.), and (c) geochemical sampling for evidence of anornalous concentrations of Hg, TI, Sb, Ba, Au, and Ag. Rapid and inexpensive methods may no t be available for a l l diagnostic cr i ter ia , and t h e risk of proceeding without them ( the other diagnostic cr i ter ia may suffice) will have t o be weighed agains t more expensive and time-consuming methods. Once a sui te of exploration methods has been matched t o t h e model's diagnostic cr i ter ia , s t r a t eg ic exploration opportunities a r e identified and, finally, t he risk of applying t h e model t o t h e proposed exploration a reas is assessed (see S t r a t eg ic Factors). If t he re a r e specific reasons why t h e program might be successful where o the r s have failed or failed t o try, exploration should proceed.

Application of a DPC model t o exploration commences a t whatever scale i s d i c t a t ed by previous exploration and cu r ren t competit ion (see S t r a t eg ic Factors). In regions relatively unexplored for a deposit type, da t a for regional and generally inexpensive diagnostic c r i t e r i a a r e collected f i rs t , a s a screening tool. If processes fo r these c r i t e r i a can be confirmed, for example, volcanism and plutonism in t h e case of

hot-spring deposits, exploration proceeds t o more detailed observations. Since volcanic-plutonic regions and t h e hydrothermal sys tems they contain a r e reasonably well known in t h e United Sta tes , s t r a t eg ic exploration opportunities for hot-spring deposits will most likely be on t h e more detailed field reconnaissance scale. Reconnaissance and detailed mapping fo r evidence of sinter, t he various hydrothermal eruption breccias, stockwork, and multiple s t ages of e a c h will de t e rmine t h e likelihood t h a t t h e episodic hydrothermal eruption process opera ted in rocks c lose t o t h e surface (a requirement for an open-pit deposit). The mapping method will also determine t h e likelihood of episodic hydrothermal fluid flow by collecting d a t a fo r diagnostic c r i t e r i a t h a t indicate t h e e x t e n t and number of s tages of a l tera t ion and mineralization.

Exploration a t t e m p t s t o identify exploration a reas in which d a t a fo r diagnostic c r i t e r i a indicate t h a t all format ion processes of a deposit type operated. Such a reas a r e highly prospective for t he occurrence of a deposit and warrant detailed exploration by more expensive methods, such a s trenching, drilling, geophysical surveys, and so forth. In t h e case of t h e hot-spring deposit type, an a r e a with evidence fo r some combination of s in ter or e j ec t a blankets, episodic brecciation, episodic hydrothermal fluid flow, and meta l l iza t ion (anomalous concentra t ions of some combination of Au, Hg, TI, Sb, and Ag) is highly prospective. Conversely, d a t a t h a t reliably document t h e absence of a format ion process condemn t h e a r e a fo r occurrences of t h e deposit t ype a s i t i s presently defined. Deposits a r e formed by chains of geologic processes, e a c h of which is essential t o deposit formation. Regardless of how extensively several format ion processes may have operated, for example, volcanism, plutonism, regional f rac tur ing and faulting, hydrothermal a l tera t ion, mineralization, and metall ization, t h e absence of one format ion process makes t h e occurrence of a deposit unlikely. For example, t h e absence of near-surface episodic hydrothermal brecciation f rom t h e preceeding process sequence makes t h e occurrence of a hot-spring-type deposit unlikely. Obviously other deposit types may be present in a reas t h a t have experienced these processes, and one may wish t o explore fo r them, but t h e potent ia l i s low fo r t h e original t a r g e t type. I t should b e noted t h a t changing deposit types in the midst of a n exploration program generally enta i ls significantly d i f ferent economic character is t ics t h a t may or may no t m e e t t h e organization's objectives. Changing deposit types i s also a human technique for avoiding loss (see Human Factors).

Experience with various mineral-deposit models indicates t h a t i t is generally eas ier t o identify high- necessity c r i t e r i a f o r a format ion process than high- sufficiency cr i ter ia . This is because many geologic character is t ics a r e always formed by a particular process ( therefore they a r e highly necessary), but t h e character is t ics a r e e i ther also formed by other processes or they do no t document a l l c r i t ica l aspects of t h e process. Fo r example, a n outcrop of quar tz- adularia mineralization documents hydrothermal fluid flow, but i t does not necessarily document the episodic fluid flow associa ted with episodic hydrothermal

explosion t h a t is apparent ly necessary for hot-spring- type deposits. The outcrop, therefore, may represent a deeper bonanza vein. Therefore , quartz-adularia is of only in t e rmed ia t e sufficiency for t h e episodic hydrothermal fluid-flow process and must be combined with o the r diagnostic cr i ter ia , such a s multiple banded mineralization and multiple crosscutting mineralized f rac tures , t o build a c a s e for episodic hydrothermal fluid flow. The g rea te r abundance of high-necessity c r i t e r i a t ends t o establish t h e permissiveness of an a r e a for format ion processes and deposit occurrence, but only high-sufficiency cr i ter ia or assemblages of intermediate-sufficiency c r i t e r i a can establish the likelihood o r favorabili ty fo r formation processes and occurrence of a deposit.

I t i s preferable, although not always possible, t o confirm t h e presence or absence of a process in a n a r e a with multiple diagnostic cri teria. This significantly increases t h e reliability of decisions t o continue o r abandon exploration programs. Where d a t a for key diagnostic c r i t e r i a a r e lacking, i t may be possible t o justify the cos t of collecting the d a t a by drilling, trenching, or o ther methods t h a t have a high likelihood of providing t h e necessary data. The availabil i ty of such d a t a fo r diagnostic c r i t e r i a should significantly increase o r decrease the favorabili ty of t h e area. The collection of da t a t h a t do not document a high-necessity or high-sufficiency cr i ter ion generally increases confusion without reducing uncertainty. If t he cos t and risk of additional da t a collection become too g rea t prior t o discovery or condemnation of a n area , t h e project is deferred until methods can be found t o continue exploration a t lower risk and cost.

Summary of Data-Process-Criteria Model

The e ight s teps of the data-process-criteria model begin with the definition of a deposit t ype t h a t presumably mee t s the s t r a t eg ic f ac to r s of ( I ) t he organization objectives, (2) a potentially profitable commodity, (3) t he organization's f inancial resources, and (4) t h e skills and style of t h e exploration organization. Completion of the model identifies t h e most informat ive and reliable diagnostic c r i t e r i a for exploration and evaluates t h e reliability of t he en t i r e model fo r exploration. Many aspects of t h e model have been used routinely by exploration geologists. However, completion of e a c h s t e p in t h e model a s described above is believed to have important advantages for exploration, t h e most significant of which a r e summarized below.

1. The DPC model is based upon sound geologic observations fo r analogs of t he deposit type. This d a t a base constrains and promotes t h e mos t reliable geologic interpretations.

2. The explicit wr i t ten model promotes communica- t ion with colleagues, peer review, and, thereby, t h e development of concensus state-of-the-art models.

3. The model fo rma t promotes multiple working hypotheses and consideration of all re levant d a t a in a n a t t e m p t t o co r rec t t h e human tendency t o build causal models with negligible regard for thei r validity. In this way, the DPC model promotes t h e search for accuracy and t ru th , r a the r than simply the search for explanations.

The usefulness of geologic character is t ics for exploration is determined through the es t imat ion of necessity and sufficiency of diagnostic cr i ter ia for identifying where in exploration a reas deposit format ion processes have operated. The selection of diagnostic c r i t e r i a for exploration is justified and documented by the linking of cr i ter ia t o one o r more deposit format ion process(es) f rom which t h e c r i t e r i a a re in terpre ted t o have formed. The model identifies the minimum diagnostic cr i ter ia (i.e., one or more fo r e a c h format ion process) for which exploration d a t a will be required, in order t o assess t h e favorabili ty of an a r e a fo r a deposit. Exploration mus t include the highest necessity and highest sufficiency cr i ter ia possible for e a c h format ion process. The model approach identifies deposit types for which only empirical or low-level models can be developed with existing d a t a and concepts. The evaluation of multiple data-process links and t h e selection of multiple diagnostic cr i ter ia permit selection of the most informative, reliable, and leas t costly c r i t e r i a for exploration. Prepara t ion of a DPC model identifies where research leading t o new d a t a and concepts is most likely t o yield new diagnostic c r i t e r i a for exploration. The model approach helps avoid collection of geologic d a t a t h a t do not contr ibute significantly t o evaluation of favorabili ty fo r a deposit in exploration areas. The approach promotes the cr i t ica l evaluation of t h e model itself and es t imat ion of t h e reliability with which i t can be used in exploration.

CONCLUSIONS

Geologic information, in the form of d a t a and concepts, plays an integral and more complex role in s t r a t eg ie s for mineral exploration than is generally appreciated. In addition t o i t s obvious use in se lec t ing and evaluating exploration areas, geologic information significantly influences (1) selection of organization objectives, (2) selection of commodities for exploration, (3) es t imat ion of financial resources required for successful exploration programs, (4) organization skills required for particular types of exploration and the deposit types best sui ted to particular exploration organizations, (5) regulatory burdens t h a t particular exploration programs and mining operations can to lera te , (6 ) evaluation of compet i tor activit ies, (7) previous exploration on proposed exploration lands, (8) selection of exploration methods, (9) development of exploration opportunities, and (10) assessment of exploration risks. The contribution of geologic information t o t h e evaluation of e a c h of these s t r a t eg ic f ac to r s is a s cr i t ica l t o t h e development of successful exploration s t r a t egy a s i t is t o t h e more obvious geologic exploration functions.

Consideration of human fac to r s indicates t h a t exploration will benef i t f rom modeling techniques t h a t help geologists control and guide cer ta in human behaviors and tendencies in thei r use of geologic observations and interpretations. I t has been

S. S. ADAMS 297

demonstra ted t h a t mos t humans instinctively (1) build causal models t o explain relations among observations, (2) t end t o disregard t h e validity of these models, and (3) resist correct ing t h e models even when new d a t a clearly indicate t h e models a r e in error. These behaviors appear t o be too fundamental t o human na tu re t o expec t much change in people's basic behavioral patterns. However, if geologists a r e aware of these behaviors and use modeling techniques t h a t minimize thei r damaging influences, t he use of geologic information in exploration should benefit. The data-process-criteria model is intended t o be a s t e p in th is direction.

Finally, t he data-process-criteria model promotes t h e more rigorous and reliable use of geologic information by making explicit (1) t he deposit type, (2) i t s analogs, (3) geologic character is t ics of t h e analogs, (4) processes responsible for t he geologic character is t ics and t h e deposit type, (5) t h e most reliable and informative cr i ter ia for exploration, and (6) t h e re la t ive importance of each criterion for evaluat ing exploration favorability. The explicitness of DPC models promotes t h e crit ique and improvement of t h e model by the model author and colleagues. Multiple working hypotheses a r e promoted by a ser ies of techniques a t every s t e p in model development in an a t t e m p t t o develop models t h a t a r e accura t e ra ther than simply explanatory. I t is hoped t h a t t h e model fo rma t n o t only promotes t h e natura l human tendency t o develop causal (i.e., process) models, but t o discontinue thei r development in favor of empir ica l models when limited d a t a and concepts warrant. Emphasis on model evaluation should help def ine t h e level of model development t h a t can be supported by existing d a t a and concepts and demonstra te t h a t inadequacies in the model a r e due to l imitations in d a t a and concepts, not in the geologist's skill. This, in turn, may make i t eas ier for geologists t o a c c e p t t h e l imitations of their models and accep t correct ions when new geologic information warrants.

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Length ~ t ~ c h e s , in fret, ft yards, yds rl'ltute miles, mi fathoms .~ngstroms, A A

Area ln2 f t 2 yds2 m12 acres 'tcres

Volume (wet and dry) in' f t ' yds3 fluid ounces quarts U.S. gallons, gal U.S. gal acre-ft barrels (oil), bbl

Weight, mass ounces avoirdupois, avdp troy ounces, oz pounds, Ib long tons short tons oz mt-'

Velocity f t sec-' (= ftlsec) mi hr- ' mi hr- '

Selected conversion factors"

Pressure, stress Ib in-' ( = lb/in2), psi Ib in-' Ib in - 2

atrn atm inches of Hg (at 0" C) bars, b b b b

Density Ib in-3 (= lb/in3)

Viscosity poises

Discharge U.S. gal min-', gpm gPm ft3 sec-'

Hydraulic conductivity U.S. gal day-' ft-2

Permeability darcies

Transmissivity U.S. gal day-' ft-' U.S. gal min-' ft-'

Magnetic field intensity gausses

Energy, heat British thermal units, BTU BTU BTU Ib-'

Temperature "C + 273 "C + 17.78 "F - 32

I < I < k>NVERT MULTIPLY BY TO OBTAIN

kg ( = kg/cm2) atmospheres, atm newtons (N)/m2, N m-2 kg ~ m - ~ mm of Hg (at 0' C) kg ~ m - ~ kg cm-> dynes ~ m - ~ atm megapascals, MPa

TO CONVERT MULTIPLY BY TO OBTAIN

centimeters, cm meters, m m kilometers, km m cm micrometers, Fm

cm2 m2 m2 km2 m2 hectares, ha gr cm-' sec-' or dynes ~ m - ~

cm3 m3 m3 liters, I or L 1 1 m3 m3 m3

grams, gr gr kilograms, kg metric tons, mt mt parts per million, ppm

gammas

calories, cal kilogram-meters, kgm cal kg-'

"K (Kelvin) "F (Fahrenheit) "C (Celsius)

(see ft2) = 0.0929. 'Divide by the factor number to reverTe conversions. Exponents: for example 4.047 x 10 (see acres) = 4,047; 9.29 x 10-

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Geology and Geochemistry of Epithermal Systems - Reviews Vol 2-Search