Collaborative research proposal between CSIRO, AGCRC, University ofWestern Australia, AMIRA and mineral exploration companies

Hydrothermal Systems,

Giant Ore Deposits

&

A New Paradigm for

Predictive Mineral Exploration

A discussion paper

written by John L. Walshe

with contributions from

Bruce Hobbs, Alison Ord, Phaedra Upton, Chris Ryan and Gem Manning

Please Consider

2TABLE OF CONTENTS

PROLOGUEThe Question 4The Acknowledgments 4The Rationale 4Diversity and complexity 5Integration and scale 5

SOME HYPOTHESES REGARDING THE FORMATION OF GIANT HYDROTHERMAL ORE DEPOSITS1. The super-giant to behemothian porphyry Cu-Au deposits 62. Giant and high grade lode Au deposits 83. The origins of the super giant Bingham Cu-Au deposit 11Crustal-scale fractures and distribution of major deposits 12The Issues 13

PART A - MAGMATIC HYDROTHERMAL SYSTEMSA1. Diversity and complexity 16A2. Synopsis of the proposed subsystems 18A3. Ortho-magmatic versus para-magmatic subsystems 19A4. The evolutionary stages of porphyry systems 20A5. Insights from the fluid inclusions 23A6. Summary correlation of vein paragenesis, alteration mineralogy and fluid inclusion data 32A7. Wriggly veins versus the straight-walled veins 34

PART B - UNDERSTANDING THE ROLE OF THE SUBVOLCANIC CHAMBERB1. Some hypotheses about barren versus productive systems 36B2. The Goonumbla Volcanic Complex 37

PART C - HYDROTHERMAL SYSTEMS IN SEDIMENTARY BASINS ANDMETAMORPHIC TERRANESC1. Key Issues 40C2. Low Temperature Hydrothermal Systems 40C3. Mississippi Valley Pb-Zn deposits - A useful analogue for low temperature systems 45C4. Some constraints on the Proterozoic Pb-Zn deposits of Northern Australia 52C5. The Mount Isa Copper deposit 53

PART D - CRUSTAL SCALE SYSTEMS, THE THIRD DIMENSION AND LODE GOLDDEPOSITSD1. Key issues 66D2. The third dimension 66D3. The Ashanti Au belt, Ghana, west Africa 68

PART E - XENO-MAGMATIC HYDROTHERMAL SYSTEMSE1. The origin of the Porgera Au deposit 76E2. The origin of the Kidston Au deposit 79E3. Bingham 87E4. The Golden Mile 93

3PART F - NUMERICAL MODELLINGF1. The Goals 99F2. A New Exploration Paradigm 100F3. Background to Quantitative Hydrothermal Modelling 101F4. Coupled Mechanical/Fluid Flow Numerical Modelling 104F5. Thermal and Deformational Modelling of the Yilgarn Deep Seismic Transect 107

PART G - FLUID COMPOSITIONS via INCLUSION ANALYSISG1. PIXE Microanalysis of Fluid Inclusions 116G2. Vapour Transport of Ore Metals 119G3. Fluid Characterization of Giant Orebodies 120G4. Fluid Composition as a Fingerprint of Fluid Source 121

EPILOGUE 123

REFERENCES 124

4Prologue

The QuestionHow and where does nature create giant, high grade mineral deposits?

The AcknowledgmentsChris Heinrich and Bruce Hobbs are to blame for the excessively enthusiastic title.

The RationaleThis proposal reflects the increasing consciousness within the community that although there is

now a general understanding of the processes which lead to deposit formation there is a lack of

understanding of the specific processes or combination of processes which are required to form

large tonnage, high grade deposits of a particular commodity or commodities.

It also reflects a growing awareness that regional-scale data, particularly geophysical data, are

fundamentally modifying perceptions of the physical dimensions of hydrothermal systems. It is

increasingly possible to think in terms of ore-forming systems rather than ore deposits. Ore genesis

theories and exploration concepts need to be re-evaluated in the light of such data.

Alan Clark (1993) has conceded that it not possible to develop a set of descriptive criteria foroutsized mineral deposits, at least for the porphyry clan. This proposed research will address the

fundamental issue of why giant ore deposits by a judicious combination of the empirical andconceptual approaches to understanding ore-forming processes. Two themes woven into the fabric

of this synopsis may well be germane to genuine progress in understanding the origins of the giant

deposits.

These are:

- the need to recognize the diversity and complexity of the processes- the need to integrate all of the data from the microscale to the crustal scale

5Diversity and complexity

The history of ore genesis theory over several centuries has been dominated by two competing

ideas as to the origins of ore fluids:

Ore fluids are derived by fluid-rock reactions.

Ore fluids are segregates from silicate melts.

A related issue which has also dominated ore genesis studies is the relative importance of sources

of metals, or provenance, versus transport and depositional processes. A major challenge in thedevelopment of robust models of the ore forming processes is the resolution of these long standing

questions to provide a general theory of hydrothermal systems. Appreciating the diversity and

complexity of the processes which lead to the formation of hydrothermal ore deposits will aid this

process immeasurably.

Integration and scale

An enhanced understanding of the relative roles of provenance and process will require a conscious

effort to integrate the microscale and mesoscale data (which provides the understanding of process)with the district to crustal-scale data (which provides the perspective on the scale of hydrothermalsystems and on provenance). Defining the size of the overall system, which commonly will bemany orders of magnitude larger than the size of the deposit or group of deposits, may also define

the size of a prospective region. For hydrothermal magmatic deposits the size of the fluid reservoir

- the subvolcanic chamber - may well define the size of the system. Alternatively, it may be the size

of a pressure-seal which operated at the district scale, which defines the size of the system.

6Some Hypotheses Regarding the Formation of Giant Hydrothermal OreDepositsThese are thumb-nail sketches only. The hypotheses are discussed in more detail in later sections of the paper.

1. The super-giant to behemothian porphyry Cu-Au depositsEXAMPLES : Chuquicamata

El Teniente

Grasberg

Breccia pipe

Two phase region, stage 2

Potassic zone, stage 1

0

2

4

6

Km

Regional seal, s tage 2

Zone of leaching, stage 2

Ore zone, stage 2

Figure P.1

7Commonly these deposits form by a two stage process (Zentilli et al, 1995)1. Formation of high temperature assemblages : M veins

: A/B veins

2. Remobilizing and upgrading of this proto-ore by lower temperature fluids to form

quartz-poor sulfide veins (C veins).The bulk of the resources form during the second stage.

Keys to deposit formation are: Development of a sealed magma chamber which acts as a fluid reservoir.

Focused release of the fluids (brine, vapor, melt) along some major conduit and commonly oldbasement structures utilized.

High fluid pressures during stage 1 to maintain sulfide saturation at high temperatures.

Condensation of acid volatiles, coupled with a drop in fluid pressure drives the leach process in

the deep parts of the system, during stage 2.

Same structures utilized during stages 1 and 2.

Quiescent phase of volcanism. A decrease in fluid pressure between stages 1 and 2 maybe related to rapid uplift.

Limited movement on the structures which act as major conduits coupled with limiteddeformation of fluid reservoirs.

Relatively stable tectonic setting, during stages 1 and 2.

In the Northern Chile porphyry province these conditions have been sustained along the west fissure

and maintained for approximately 10 Ma from the Late Eocene to Early Oligocene. This system

was most probably driven by a semi-continuous magma chamber beneath the whole belt. Domains

of fluid focusing were controlled by crosscutting NE-trending basement structures.

8

102 1031 10 km10

100

5

10

km

1

Region of melting

GroundwaterReservoir

Two phase region

Brine & gas reservoir within

magma chamber

2. Giant and high grade lode Au deposits EXAMPLES : Ashanti

Golden Mile

Porgera

Lihir?

Telfer

Northern Nevada Au province

These deposits developed at many different times in earth history and in a range of geological

settings. They:

May be associated temporally and/or spatially with intrusive rocks.

May be hosted by sequences of black shales.

Commonly occur within large fault systems or related to crustal lineaments.

Commonly evidence of hypogene remobilization of gold and release of refractory gold.

Figure P.2

9 Appear to have formed a various levels in the crust. A common theme is formation at a late

stage in a tectonic cycle.

Fluids commonly CO2 and/or CH4 rich aqueous fluids. Some are notably depleted in H2O and

may be nitrogen rich.

Commonly evidence of strong redox contrasts at the site of ore deposition. In such cases Au

appears to have been transported by the reduced fluid and deposited by oxidation.

It is suggested that:

Au is mobilized from deep-crust/mantle in the upper crust by variously reduced volatile-rich

fluids. The fluids are of composite origins and may include surface/high level fluids that

recharge aquifers in the lower crust during periods of extension.

Depositional mechanisms are variable but a significant process is the mixing of a reduced Au-

bearing fluid with an oxidized fluid.

In the Ashanti deposit the redox contrast was generated by mixing a deep-crustal CH4-N2- rich fluid

with a CO2 -rich fluid at mid-crustal levels. Significant hyogene reworking also occurred.

Zone 7 of the Porgera deposit resulted from the condensation of SO2-rich magmatic volatiles into a

reduced fluid within the Romane Fault. A significant proportion of the Au was probably reworked

from the earlier A-vein veins (Cameron et al., 1997)

The Golden Mile developed at a high level in the crust by mixing of a reduced, gold-bearing CO2-

rich fluid (the generic Archean gold fluid of Phillips et al., 1996 that deposited the thousands ofother Archean lode deposits throughout the Yilgarn) with a SO2-rich magmatic fluid. The chemicalintegrity of the two fluids up to the point of mixing was maintained by the Golden Mile Dolerite

which behaved as an aquiclud at the local to regional scale.

At Kidston, the Lochaber - Bagstowe complex was a reservoir of mildly oxidized, CO2-rich gases

and the Gilberton structural corridor acted as a flow zone for the reduced fluid. At the regional

scale, the Kidston pipe, was a domain of low pressure drawing in fluid from both the Gilberton

corridor and the Lochaber - Bagstowe complex. The pipe effectively operated as a valve for both

fluid reservoirs.

10

Lower part of sequenceGolden Mile DoleritePervasive Carbonate Alterat ion

LodesShears

A B

400 m

600 m

1000 m

800 m

200 m

200 m

Reduced Au - r ich Fluid

Mixing within the lodesof reduced Au f luid andoxidised magmatic f luid

Black Flag Beds

Key to size and grade:

Regional scale focusing of fluids from two reservoirs into a domain of low pressure.

This may occur at various levels in the crust.

The bonanza grades reflect the strong redox gradients generated by simultaneously focusing two

fluids of strongly contrasting redox conditions.

3. The origins of the super-giant Bingham Cu-Au deposit

Figure P.3

11

0 300 km

Carlin Trend IndependenceTrend

GetchellTrend

Bingham Canyon

Alligator RidgeTrend

Battle Mountain Eureka Trend

Sevier orogenicbelt

Battle MountainDistrictCopper BasinCopper CanyonFortitiude

Trend of 35-43 Ma Intrusions

The Key: Regional scale focusing of two fluids?? The size and Au tenor of the Bingham Cu-Au deposit is unique in the southwestern US

porphyry province.

It is hypothesized that this super-giant porphyry deposit resulted from the effective focusing and

mixing of two discrete fluids : a Cu-rich but sulfur-depleted magmatic brine and

: a reduced, H2S-rich and Au-rich fluid.

The later fluid is akin to the fluids that formed the sediment-hosted Au deposits of northern

Nevada.

Figure P.4

12

The productive period of Cu-Au deposit formation within the Great Basin reflects this

interaction between magmatic fluids and deep-seated reduced Au-bearing fluids at around 40

Ma across Northern Utah and Nevada.

Crustal-scale fractures and distribution of major deposits

Empirical evidence supports an association of significant deposits and mineral provinces with

major faults and/or lineaments.

The origins of lineaments remain controversial, particularly as such features appear in part to be

independent of upper crustal geology. Here, it is assumed that they image an array of fractures

within the mantle and that these mantle fractures have repeatedly influenced patterns of

deformation in the crust.

A first-order set of mantle fractures is estimated from the outline of the Australian continent. The

rather geometric shape of the western two-thirds of the continental is taken to reflect the influence

of an array of mantle fractures on the break-up of Gondwanaland during the Mesozoic and

Cenozoic.

The fracture sets also played a significant role during phases of convergence. The active set

appears to relate to the angle of collision.

The north-east trending set was active during collision of the Australian and Pacific plates during

the last 10 ma, acting as transfer structures and controlling the distribution of major mineraldeposits in Papua New Guinea and Irian Jaya (Hill and Mason, 1997; Hall, 1995; Corbett, 1994).

The west-northwest trending set exerted a fundamental control on the distribution of majordeposits and mineral provinces in the Tasman Orogen during Paleozoic convergence events

(Walshe et al., 1995; Walshe and Glen, 1996). Greg Hall (1996) has suggested that the golddeposits in Kalimantan relate to a north-northwest set of lineaments. This trend correlates with a

similar trend in the Australian continent.

13

45o

Lach lan RiverTransfer Zone

OrdovicianLarapintine Sea

NorthQueenslandAu prov ince

Centra l-west NSWCu-Au prov ince

PNG - Irian JayaCu-Au province

KalimantanAu prov ince

The Issues Fluid reservoirs Fluid mixing, particularly liquids and gases Hypogene enrichment of primary mineralization Controlling basement structures Pressure evolution with depth, temperature and time

The construction of giant, high grade hydrothermal deposits requires effective methods of storing

large volumes of fluids within the earths crust while sustaining metal and sulfur solubility within

Figure P.5

14

those fluids. There must also be effective ways of delivering the fluids to sites of ore formation,

whilst ensuring the chemical gradients necessary to produce high grade mineral deposits are

maintained.

Metal solubility is essentially controlled by a relatively small number of parameters: temperature,

acidity of fluid, redox state of fluid, salinity of fluid and concentration of volatile species,

particularly sulfur. A fundamental constraint on the development of most base metal deposits is the

need to deliver sufficient reduced sulfur, as well as, metals to the site of ore formation.

High metal grades will occur when large gradients in one or more of these solution parameters can

be sustained. The mechanisms for sustaining these gradients are limited. Common processes

include reactions of fluids with specific host rocks, phase separation at particular sites and mixing

of fluids (liquids and/or gases). Fluid-rock reactions and phase-separation mechanisms have limitedcapacity to maintain gradients and generate both large tonnages and high grades (Heinrich, 1996).Processes involving the mixing of large volumes of fluid with strongly contrasting properties or

processes involving the recycling and upgrading of initially low grade deposits are potentially the

most effective mechanisms for generating both high grade and large tonnage deposits.

It is important to understand the processes operating at the regional to crustal scales that give rise to

reservoirs of heat, salinity, acidity, sulfur and redox capacity within the crust and mantle. These are

effectively the metal reservoirs or the reservoirs of the critical reagents (oxidant, reductant, etc.)required for metal precipitation. Knowledge of the mechanisms of fluid release from these

reservoirs and the pathways which allowed fluids to be focused into the upper crust while

maintaining their chemical integrity is also significant. It is at this point that the bulk

composition/lithology of rock sequences in the upper crust are likely to play a major role inenhancing or degrading the ore forming potential of fluids generated in the lower or middle crust.

Maintaining fluid pressure as well as temperature within fluid reservoirs is likely to be of

paramount importance. Commonly, the partial pressure of the acid volatile species (CO2, H2S, SO2,HCl) determines acidity and oxidation state; sulfur levels and maintains a balance of both metal andthe sulfur components within the fluid. Loss of fluid pressure is likely to lead to a degradation of

the reservoir both in terms of volume of available fluid and its chemical potential for deposit

formation. However, in a pre-existing ore zone a loss of fluid pressure can lead to sulfide

dissolution and a significant upgrading of the resource.

15

There is a need to recognize the existence and extent of regional seals, be it a plug in the top of a

magma chamber or a clay horizon within a sedimentary basin, and to learn to recognize when one or

several reservoirs have released fluids in a controlled and focused way. These are likely to be times

of great potential for large tonnage and high grade hydrothermal deposits.

16

Part AMostly about Porphyry Deposits

Magmatic Hydrothermal Systems

A1: Diversity and complexity

It is suggested that the diversity of processes within the clan of magmatic hydrothermal depositsmay be represented in terms of the following end-member subsystems:

Magmat ic - Hydrothermal Subsys tems

1 432

1. Ortho-magmatic subsystems: - straight magmatic fluids; metal budget of the fluid is a product

of supra-solidus processes in the magma chamber.

Figure A.1

17

2. Para-magmatic subsystems: - evolved magmatic fluids which have been substantially modified

by fluid-rock reactions and/or mixing with external fluids. Metals are largely derived from the

intrusive complex but metal budgets reflect the extensive subsolidus history of these fluids.

3. Supra-magmatic subsystems: - dominantly external fluids and metal sources; the magma

chamber is a source of heat and possibly acid-volatiles, oxidant and/or some key metals.

4. Xeno-magmatic subsystems: - metal-bearing fluids sourced from similar depths to the magmas;

interplay of magmatic and xeno-hydrothermal fluids may be critical to deposit formation in the

upper crust.

These model systems summarize the dominant processes by which metals may be transported and

deposited in fluids, in and around intrusive complexes at about the time of complex formation.

These subdivisions do not replace the well defined and accepted classes of deposits which are

based on descriptive criteria. They are based on the recognition that the processes operating in and

around an intrusive complex can be broadly grouped. To some extent these subsystems may be

considered substages or subdomains since subdivisions 1, 2 and 3 could be considered stages or

domains within a single evolving system. However, there is no requirement that all of these

subsystems be present in a given system just as there is no requirement that an epithermal depositshould sit above a porphyry deposit. Conversely, few deposits are likely to have formed through

operation of just one of the dominant processes summarized by these subsystems. The genesis ofmost deposits will reflect some complex interplay between these end-member systems.

Systems are commonly open; open at the top, the sides and the bottom. We have a good

understanding of the openness of these systems at the top but a very poor understanding of their

openness at the bottom. Model 4 addresses this issue and permits consideration of the interaction

of fluids from a magma chamber with fluids from non-magmatic reservoirs that might reside in the

middle or lower crust.

18

A2: Synopsis of the proposed subsystems

The ortho-magmatic subsystem: The ortho-magmatic hydrothermal fluid is a directproduct of the silicate melt. It seems that it is possible to recognize deposits which formed from

fluids that contained a dominant component of this fluid. The best examples are found amongst

the Cu and Au rich porphyry deposits. Typically the fluids are high T; oxidized, super saline

and sodic at least in their early stages. Such deposits are commonly associated with an

underlying magmatic chamber which supplied the juice. The Burnham view (Burnham, 1979)of how porphyry deposits form probably applies more to the chamber than the overlying deposit.

In these systems the nature of the silicate melt is extremely important. Another key factor seems

to be the ability of the high level chamber to release stored fluids in a focused manner. A

critical problem is how to recognize when and how sub-volcanic chambers act in this manner.

An understanding of these processes in particular should yield more robust conceptual models

The para-magmatic subsystem: D and O isotope studies of porphyry Cu deposits andSn deposits through the 1970s and 80s brought recognition that the phyllic alteration in many

deposits reflects an ingress of meteoric water into the magmatic hydrothermal system,

commonly at a late stage. This led to a two-fluids model of these systems - a magmatic ore fluid

which essentially formed the deposit and a late-stage fluid of meteoric origin which variously

overprinted and modified the pre-existing sulfide deposit as the magmatic hydrothermal system

waned. What is less well recognized is the ore fluid which is transitional between ortho-

magmatic fluids (as defined above) and external (near surface) fluids which are readilyrecognizable from their D and O isotopic signatures. These para-magmatic fluids may be

evolved ortho-magmatic fluids which derived their metal budgets through re-equilibration with

the host igneous body at subsolidus conditions; they may be external fluids which re-

equilibrated with igneous host rocks or perhaps most commonly they are fluids derived by a

combination of these processes. The term para-magmatic is adopted here to describe these

fluids of diverse origin but which retain some definable link with the magmatic system; either

by direct derivation of some components from the silicate melts or through exchange of

components with the crystallized igneous body at subsolidus conditions. The prefixes ortho-

and para- contrast the idea of the straight magmatic fluid with that of an evolved magmatic

fluid which still retains some of its magmatic heritage. Making the distinction between these

two fluid types is important to our understanding of the complex array of processes involved in

19

the transition from the purely magmatic domain of the magmatic hydrothermal system to the

open domain of the system where external fluids play a highly significant role. It is within this

transitional domain that ore fluids may be generated and dispersed; metals concentrated and

diluted; ore bodies won and lost. Many porphyry Cu deposits may have formed entirely from

para-magmatic fluids which had their metal and sulfur budgets determined by subsolidus

processes within the intrusions.

Supra-magmatic systems: Epithermal, mesothermal and other deposit types which may berelated to underlying intrusions (known or inferred). The ore fluid is largely an external fluidwith the metals being largely derived from the country rocks but with key components (acidvolatiles and some metals) being derived from the underlying intrusion.

Xeno-magmatic systems: Encompasses the concept that a hydrothermal fluid may comefrom a deep-seated source(s) in the crust and/or mantle and may ascend similar pathways as themagmas. The fluids would need to be CO2 or CH4-rich at depth so as not to dissolve in the

melts and may be highly reduced. Mineral deposition may occur through mixing with surface or

magmatic fluids, or through reaction with specific host rocks. Chances are we are not very good

at differentiating 4 from 1, 2 and 3 in high-level geological settings. Deep-seated reduced fluids

which find their way into the upper crust may generate reduced Au-skarn deposits for example.

Such deposits are commonly spatially associated with oxidized magmatic systems (e.g. SheahanGrants, Fortitude, Ladolam) and the switch from high-T oxidized fluids to low-T reduced fluidsseen in these systems could be explained by a switch from oxidized magmatic fluids to deep-

seated reduced fluids.

A3: Ortho-magmatic versus para-magmatic subsystems

In brief It seems possible to make a first-order distinction between porphyry deposits which formed by

ortho-magmatic processes in the main and those deposits which formed by para-magmatic processes

from the nature of veining, the alteration assemblages and the fluid inclusion compositions.

20

The salient characteristics of deposits in which ortho-magmatic processes played a dominant role

are:

Early high temperature Fe - Na Ca metasomatism characterized by plagioclase biotite and

magnetite

This early assemblage is commonly overprinted by a potassic assemblage of K-feldspar - biotite

magnetite - sulfide anhydrite.

Irregular, discontinuous quartz veins which predate laminated straight-walled veins.

Highly saline (> 50 wt. %) to supersaline fluid inclusions (up to 70 to 90 wt %); commonly themost saline fluids are the most sodic.

Experimental studies and thermodynamic calculations indicate that the high salinity, sodic fluids

were most likely in equilibrium with a silicate melt at suprasolidus conditions, prior to the

saturation of alkali feldspar and quartz, at high temperatures and low pressure.

The porphyry deposits most likely to have formed by para-magmatic processes are those in which

the deposition of sulfides is relatively late in the paragenesis. In these deposits an early sulfide-

poor quartz stockwork is commonly associated with potassic alteration (K-feldspar - biotite).Sulfides are deposited from low salinity fluids at relatively low temperatures.

A4: The evolutionary stages of porphyry systems:

Beginning at the mesoscale: vein paragenesis and alteration assemblages: The evolution of porphyry systems can be described in terms of the generalized paragenetic

scheme:

- early M and A veins

- transitional B veins

- late C and D veins

This is simple retrogressive paragenetic scheme.

21

The histories of deposits show infinite complexity but most (all?) seem to be some variation onthis basic scheme.

Vein Paragenesis

Understanding the complex interplay between spatial relationships, highlighted by Lowell and

Guilbert (1970) and the temporal evolution, emphasized by Gustafson and Hunt (1975) in theirstudy of the El Salvador deposit, remains fundamental to unraveling the histories of magmatic-

hydrothermal systems. Gustafson and Hunt (1975) used a scheme of A, B and D veins and relatedalteration assemblages which reflected the temporal evolution observed in that deposit. The scheme

has been extended by Dilles and Einaudi (1992) to include C veins and by Clark (1993) and Clarkand Arancibia (1995) to include M veins to give an overall evolutionary scheme for porphyrysystems of M veins (commonly earliest) through A, B, C and D veins (commonly latest).

This generalized paragenesis, represents the evolution from early high-temperature alteration

processes (M and A stages) to later lower temperature processes (C and D stages) with Brepresenting a transitional stage. The significance of the M stage has only recently been recognized.

It may be that in many instances it has been overlooked as a hydrothermal event or confused with

other sodic and calcic alteration events which may develop as cooler fluids on the margins of

systems are drawn in and heated (Carten, 1986, Dilles et al., 1992, Dilles and Einaudi, 1992). Few deposits exhibit the full spectrum of vein types and alteration assemblages.

Examples of deposits with early M and/or A stages: Endeavour 26N, Goonumbla Volcanic Complex, NSW

Island Copper, Cu-Mo-Au deposit, British Columbia

Park Premier, Central Wasatch Mountains, Utah

Yerington, Yerington Batholith, Nevada

El Salvador, Atacama Province, Chile

Grasberg, Grasberg Igneous Complex, Irian Jaya

Koloula Porphyry Copper Prospect, Guadalcanal

Panguna Cu deposit, Bougainville, Papua New Guinea

Bajo de La Alumbrera, Farallon-Negro Complex, Catamarca, Argentina

22

High temperature and commonly early

M veins

Magnetite or quartz-magnetite veins associated with sodic calcic alteration assemblage of

plagioclase (andesine through albite) biotite calcic amphibole (actinolite to magnesio -hornblende) pyroxene.

A veins

Irregular, discontinuous veins of quartz magnetite bornite chalcopyrite anhydrite

associated with a potassic alteration assemblage of K-feldspar biotite.

B Veins (transitional)

Straight-walled, continuous, laminated, gray vitreous quartz veins, centimeters thick

commonly with molybdenite chalcopyrite bornite magnetite anhydrite. Sericite may

develop in the potassic alteration assemblages in this stage as well as tourmaline veins.

C veins

Veins of quartz biotite chlorite epidote chalcopyrite bornite pyrite

molybdenite.

D veins

Veins of quartz pyrite chalcopyrite anhydrite sphalerite galena associated with an

alteration assemblage of quartz - sericite - chlorite.

Low temperature and commonly late

23

In porphyry deposits associated with shoshonitic complexes (Grasberg, Endeavour 26N) the earlystage mineralogy appears to be limited to magnetite-biotite-plagioclase whereas for deposits hosted

by calc-alkaline complexes (Koloula Porphyry Copper Prospect, Park Premier, Island Copper Cu-Mo-Au deposit) the early alteration mineralogy commonly includes amphibole and pyroxene.

In many deposits the early M and A veins are not developed and the earliest veining is a barren

quartz stock (most probably equivalent to B) with minor sulfides, most commonly molybenite.

Examples of deposits without early M and/or A stages: Ann-Mason Cu deposit, Yerington Batholith, Nevada

Bingham Cu deposit, Utah

Ok Tedi Cu-Au deposit, Western Province

Santa Rita Cu deposit, New Mexico

Sierrita Cu deposit, Arizona

Taking account of the microscale data: A5: Insights from the fluid inclusions

Summary comments In the ortho-magmatic systems fluid compositions converge with increasing temperature to

extremely saline and sodic compositions.

These compositions are to be anticipated for a saline brine coexisting with a silicate melt, a

magmatic vapor and plagioclase alkali-feldspar.

M-stage veins with their characteristic sodic calcic alteration assemblage of plagioclase

(andesine through albite) biotite calcic amphibole (actinolite to magnesio-hornblende) pyroxene are generated by ortho-magmatic fluids which coexisted with silicate melt prior to

saturation of alkali-feldspar in the melt and most probably prior to saturation of quartz.

The characteristic potassic alteration of A-vein stage occurs after alkali-feldspar saturation in

the melt. Cooling of fluids from a melt saturated in the two feldspars leads to the precipitation of

K-feldspar at the expense of plagioclase.

24

3

NaCl KCl

H2O

2

1

PangunaGranisle-BellRed MountainEndeavour 26N (Goonumbla)

??

4

4

Type C

Type D

20 40 60 80

NaCl (weight percent)

Compositions of the saline fluid inclusions

Fluid inclusion studies of magmatic hydrothermal deposits have shown that the total salinity of the

hydrothermal fluids may vary from near zero through to eighty or ninety weight percent NaCl

equivalent. The high salinity inclusions are generally taken to be magmatic brines and a first order

view, consistent with other data (geological setting and stable isotopes), is that variation in fluidsalinity is commonly a result of dilution of high temperature magmatic brines with dilute surface or

ground waters. However other sources of salt certainly cannot be discounted nor other mechanisms

of dilution.

Inclusion compositions for porphyry deposits also show a wide range in NaCl/KCl ratio (data for aselection of deposits are shown in Figure A.2). An important feature noted from numerous studies

Figure A.2

25

of porphyry deposits is the halite trend of Cloke and Kesler (1979) - linear to sub-linear trends inthe data that converge towards the NaCl corner of the NaCl - KCl - H2O ternary (Eastoe, 1978,1982; Cloke and Kesler, 1979; Wilson et al., 1980; Roedder, 1984; Quan et al., 1987) withincreasing temperature. The origin of the halite trends was much discussed in the literature of the

1970s but a satisfactory explanation remained elusive. The following interpretations developed

from the study of the Endeavour 26N deposit in the Goonumbla Volcanic Complex (Heithersay andWalshe, 1995).

Not all porphyry deposits show these trends. They appear to be associated with porphyry deposits

which have the early M and A stages of mineralization (Goonumbla, Panguna, Park Premier,Granisle Bell). Deposits which only contain B, C, D stages (Bingham, Santa Rita and Sierrita) donot show the halite trend; inclusion salinities in these deposits range between about 5 and 50 weight

percent NaCl equivalent.

26

The composition of the ortho-magmatic brine

Mi x i ng&

di l u t i onof

m ag m at i cbr i ne

400

600

800

200

1000

1000 20 40 60 80

Sa l in i ty (w t% N aCl )

500 bars

1000 bars

Tem

pera

ture

(oC

)

Dom a i n o f f l u i dho m oge n i z a t i on

Gr o un dwa t e r

Porphyry Cu-Au

MagmaticVapor

Ortho-magmaticBrine

The salinity of a magmatic fluid may be controlled by the extent of magma fractionation but once

the two phase region (coexisting vapor and brine) is intersected (commonly at high temperature and

Figure A.3

27

low pressure) the salinity of the magmatic brine will depend only on temperature and pressure.Hence at the P/T conditions most probably encountered in high level magma chambers (saypressures around 500-750 bars and temperatures of the order of 700-1000 C) the total salinity ofthe fluid will lie in a relatively restricted range from about 70 to 90 weight percent (Figure A.3).

Holland (1972)

0 20 40 8060

Total Sal inity (wt. percent)

0.4

0.2

0.6

0.8

1.0

Kravchuk et al. , 1993

(mK/mNa)melt

(mK/mNa)aq

If a concentrated brine coexists with a silicate melt, NaCl is strongly partitioned into the brine

relative to KCl, i.e. the ratio of (mK/mNa)aq to (mK/mNa)melt is less than 0.1 (Figure A.4 and theexperimental work of Kravchuk et al., 1993). Hence, the composition of very saline fluids,separating from a silicate melt must lie close to the NaCl - H2O side of the ternary almost

irrespective of melt composition.

These two observations suggest that the commonly observed convergence of fluid inclusioncompositions at high temperature to high salinities and sodic compositions reflects a convergenceto a common fluid composition buffered by silicate melt and magmatic vapor.

Figure A.4

28

Such fluids will precipitate plagioclase on cooling. Once the melt is saturated in two feldspars,

plagioclase and alkali feldspar, the NaCl/KCl ratio of the brine, vapor and melt will be

approximately buffered by these minerals (approximately but not uniquely because the compositionsof the two minerals also evolve). These fluids will precipitate alkali feldspar on cooling.

KC l

H2O

0.1 0.25 0.5 1

mKCl / mNaCl

2

High T, high P

Para-magmaticFluidsOrtho-magmatic

Fluids

NaCl

High T, low P

Melt+

plagioclase

Melt+

plagioclase+

alkali feldspar

Albite+

K-feldspar

K-feldspar

muscovite

Vapor+ saline fluid

in excess

Compositions of rock-buffered fluids within igneous complexes

As most igneous rocks contain two feldspars at subsolidus conditions, the conditions of fluid-rock

equilibrium are equivalent to the domain in which the two feldspars are in equilibrium with the fluid

in Figure A.5. This domain is based in part on experimental work and in part on observations from

Figure A.5

29

natural systems. At high temperatures the two feldspars will coexist with a silicate melt, brine and

vapor, this condition is shown as a separate subdomain in Figure A.5. At subsolidus conditions the

two feldspars coexist with brine and vapor only. For deposits such as Panguna it appears from the

fluid inclusion compositions that the fluids were close to equilibrium with the two feldspars

(compare Figures A.2 and A.5). For most deposits the fluid inclusion data indicates significantfluid-rock disequilibrium.

It is possible that the very sodic fluids are of straight magmatic origin. However, this is not possible

for the many fluids which fall on the more potassic side of the two feldspar domain shown in Figure

A.5. These fluids must be the products of subsolidus fluid-rock reaction and may be generated by

an acid leach (perhaps driven by acid volatiles condensing from magmatic gases) of the host rocks.Their compositions will depend on the composition and mineralogy of the host rocks as well as the

initial fluid conditions.

Ortho-magmatic versus para-magmatic saline fluids Revisiting the concept

Within porphyry deposits, it seems possible to identify saline fluids, or at least components thereof,

which are of straight magmatic origin (the ortho-magmatic component) and fluid components (thepara-magmatic component) which have undergone reaction and exchange with the host rocks and/ormixed with external fluids.

The significance of differentiating between these fluids lies in the potential controls on their metal

carrying capacity. The capacity of the para-magmatic fluids may be largely a function of their

subsolidus history while the ore forming potential of the ortho-magmatic fluids will reflect the

properties of the magma.

30

NaCl

H 2 O

KCl

High T, h igh P

Para-magmat i cFluidsOrtho-magmat icFluids

High T, low P

Indicative compositional ranges for these two fluid types on the NaCl - KCl - H2O ternary are given

in Figures A.5 and A.6.

The domain of ortho-magmatic fluids is drawn assuming:

Saline fluids of around 70 to 80 weight percent may coexist with two feldspars plus silicate

melt and vapor at high temperature and low pressure,

and that

Lower salinity ortho-magmatic fluids may coexist with plagioclase only at still higher

temperatures and pressures. These are potentially the most primitive magmatic fluids seen

in porphyry deposits.

The domain of para-magmatic fluids encompasses fluid conditions which approximate equilibrium

with two feldspars as well as fluids with higher ratios of KCl to NaCl. These higher ratios are

indicative of fluids modified by subsolidus fluid-rock reaction.

Figure A.6

31

Subsolidus reactions and para-magmatic fluids

The mechanisms for achieving values of mKCl/mNaCl greater than the ratio given by the albite-K-

feldspar buffer are limited. If K-feldspar remains stable, as is commonly the case in the inner zones

of porphyry systems, it may occur once all the albite is removed and if the zone remains closed to

external waters. In this way the mNaCl of the fluid will be constant as it is neither added nor

removed from the fluid. The mKCl content of the fluid will be increased by the dissolution of K-

feldspar in the potassic zone to form sericite and this reaction can be driven by the partial pressure

of HCl in the system. In this situation the governing fluid-rock reaction would be

3KAlSi3O8 + 2HCl v + 2H2Ofld KAl3Si3O10(OH)2 + 2KClfld + 6SiO2 + 2H2Ov

for which

(mKCl/H2O)fld = (HCl /H2O)vap . K

Hence the mKCl/mH2O of the brine will increase with increasing partial pressure of HCl in the

system. This will yield an increase in mKCl/mNaCl providing mNaCl is constant.

This argument suggests that the high KCl/NaCl and KCl/H2O ratios of the saline fluid inclusions in

deposits such as Endeavour 26N (type D), and Granisle Bell indicate extensive brine-vapor-rockinteraction at high partial pressure of HCl. It is possible that gas pressures were generally high in

these systems. The linear to sub-linear trends on the NaCl - KCl - H2O ternary for these deposits

suggests mixing of ortho-magmatic fluids with brines that evolved within the zone of potassic

alteration. Mixing of these ortho- and para-components may be an important process in sulfide

deposition.

It may be that systems which apparently evolved within the two-feldspar domain with falling

temperature and salinity were systems which lost pressure with cooling. It is likely that an influx of

external rock-buffered fluids was associated with any such pressure drop and this would also

promote evolution within the two-feldspar domain with falling temperature.

32

H2O

KCl

plagioclase +alkali feldspar

NaCl

plagioclase

A/B - vein fluidsK-fe ldspar , biot itesulf ides m agnet ite anhydr ite

quartz poorsulf ide r ich

D - vein fluids Com monly m olybdeniteB-vein equivalents

Barren quartz veins

M-vein fluidsplagioc lasemagnet itesulf idesFe-Mg s il icates

Integrating the microscale and the mesoscale data

A6: Summary correlation of vein paragenesis alteration mineralogy fluid inclusion data

Ortho-magmatic systems It is possible to correlate the fluid inclusion data with the vein paragenesis and alteration

assemblages. Deposits which have the M-vein stage (i.e. show early plagioclase alteration withmagnetite, biotite, amphibole pyroxene ) transitional to A/B include Park Premier, Panguna,Endeavour 26N, and Granisle Bell. The fluid inclusion data for these deposits also show the

characteristic convergence to sodic, saline fluids at high temperature.

It may be inferred that the saline brines which formed the M-vein stage separated from the silicate

melt prior to saturation in alkali-feldspar and most probably prior to quartz saturation also. Quartz

Figure A.7

33

phenocrysts in porphyries associated with this stage are uncommon in the deposits quoted. The

primitive brines were probably cooled and de-pressurized prior to formation of the quartz veins;

their high salinities may reflect the P/T conditions of fluid-melt equilibrium rather than P/T

conditions of trapping of the fluid in quartz.

The potassic alteration assemblage (K-feldspar biotite) associated with A and B veins developedby cooling of hydrothermal fluids separated from melt saturated in both plagioclase and alkali-

feldspar. Cooling promotes the precipitation of K-feldspar at the expense of plagioclase.

The extent of subsolidus reaction at the A/B stage within potassic zones may be largely controlled

by gas pressure. High gas pressure, or at least a high partial pressure of HCl, will enhance

reactivity. Maintaining a high partial pressure of H2S or SO2 may also be important in stabilizing

the sulfides at high temperature and in particular promoting the stability of bornite over

chalcopyrite.

Barren stockworks and para-magmatic systems

A loss of gas pressure in a magmatic hydrothermal system is likely to limit reactions of the brine

component with the rocks of the igneous complex and/or country rocks and may lead to brines

becoming undersaturated in sulfides. Systems in which the pH and the temperature are effectively

buffered by the rocks are prone to sulfide undersaturation as the activity of sulfur in the system

drops without a compensating drop in temperature or increase in pH which can occur in an open

vein.

This phenomenon results in sulfide dissolution if it occurs in a column of rock which contains a

significant amount of pre-existing sulfide. The textures indicative of sulfide dissolution are not

widely recognized but have been noted in cassiterite-sulfide deposits (Halley and Walshe, 1995) andat Porgera (Cameron and Walshe, 1996). The pre-existing sulfide in the rock column ensures thatthe fluids remain saturated in sulfide despite the loss of sulfur from the system. If a drop in gas

pressure occurs prior to sulfide saturation or in rocks that contain only minor amounts of sulfides

then the fluids may become undersaturated in sulfide. The dispersal of magmatic brine and vapor,

rather than the focused release of the two fluids, can lead to the brine becoming unsaturated in

sulfides. This phenomenon is probably quite common and largely unrecognized.

34

Brines undersaturated in sulfides may give rise to the early barren stockworks (most likelyequivalent to B) which occur in a significant number of deposits. Fluid inclusion studies atBingham, Santa Rita and Sierrita have shown that the barren stockworks which are associated with

potassic alteration contain the highest temperature (Reynolds and Beane, 1985) and highest salinityfluids (Preece and Beane, 1982) within the deposits and work at Bingham (Anderson et al, 1989)has shown these fluids are metal rich. However, sulfide precipitation is associated with later veins

which formed at lower temperature, from lower salinity fluids (equivalent to C and D veins in thegeneralized paragenetic scheme). In these deposits the key to sulfide precipitation could be thedecline in temperature or these degassed brines might be recharged by in-mixing of sulfur-rich

external fluids or by in-mixing of magmatic gases being discharged from deeper parts of the

magmatic-hydrothermal system.

The rationale for correlating barren stockworks with B (transitional) veins A7: Wriggly veins versus the straight-walled veins

Some preliminary observations and interpretations

A very general indicator of pressure regime.

Wriggly veins are indicative of fluid pressures just equal to confining pressure - probably abovehydrostatic but sublithostatic.

Straight walled veins are indicative of lower confining pressures and veins remain open.

Although they are laminated, they are not crack seal veins and therefore were opened a limited

number of times.

The transition from wriggly to straight occurs in the ortho-magmatic systems at the level of the

deposit.

In the para-magmatic system this transition occurs somewhere below the level of the deposit.

Differing vein styles may be used as a rough P/T delimiter in magmatic hydrothermal systems as

the vein style is dependent on P/T and independent of chemical reactions within the system.

35

Part B More on ortho-magmatic systems

Understanding the role of the Subvolcanic Chamber

Vapor+ Melt

+ Saline Brine

Km

5

0

2.5

THESEALED

TANK

Figure B.1

36

B1: Some hypotheses about barren versus productive systems

In the ortho-magmatic systems the high level chamber acts as a reservoir of melt, brine and

vapor.

The internal pressure of the chamber is critical. The productive chambers act effectively as

pressurized autoclaves at the time of deposit formation .

Focused release of brine and vapor at high partial pressure of sulfur will promote sulfide

saturation at high temperatures.

The deposit may be several kilometers above the actual fluid source within the chamber. Fluids

must be focused to these levels in the system.

Precipitation of bornite in preference to chalcopyrite will be favored by a high ratio of H2S and

CuCl (or some equivalent species) to HCl in the systems as illustrated by:

CuFeS2 + 2H2S + 4CuCl Cu5FeS4 + 4HCl

Open chambers, which vent to the atmosphere and lose internal pressure, will produce metal-

rich but sulfur-poor brines. To be effective mineralizing agents, these brines require

recharging in some way by subsolidus processes.

The Burnham view of stock-work formation and sulfide deposition within the chamber may

apply to low pressure systems. Large, low grade deposits are the likely products of such

systems.

37

B2: The Goonumbla Volcanic Complex

Interpreted Setting

The oldest volcanics in the complex, the Nelungaloo Volcanics, were probably deposited

unconformably on a quartz-rich turbidite basement. The fossil evidence indicates an age of around 480

Ma. The earliest volcanic products appear to have been high-K calc-alkaline lavas. A poorly

understood hiatus occurred until around 440Ma when development of the shoshonitic volcanic

complex began. The volcanic edifice was initially submarine, gradually building up to be subaerial at

later stages. It was fringed by shallow marine limestones.

-18

-20

-23

Simplif ied Structura l andGeophysical Features

E31E26N

Gr av i t yCon t ou r s( m i l l i ga l )

N

MagneticHigh

CalderaRim

A

B

0 5 Km

EndeavourLinear

MonzoniteIntrusions

-22

Figure B.2

38

A circular feature in the magnetics is taken to represent the outer limits of a central caldera (as shownin Figure B.2) which would have been about 20 km in diameter. Monzonite-monzodiorite intrusionsin the northern part of the complex were emplaced along the caldera rim. It is practicable to model the

pronounced gravity low at Goonumbla by assuming a monzodioritic chamber beneath the southern

part of the central caldera (Jones, 1985). Subsidence above the chamber occurred prior to theemplacement of the Wombin volcanics.

Geological In terpretat ion

present surface

A B

Cumulate

E31

E26N 2

3

4

5 Km

Geochemical patterns in the whole-rock data

Two trends are evident in the whole-rock geochemical data of Heithersay and Walshe (1995). Theyare thought to reflect fractionation within the subvolcanic chamber as magmas ponded near the base of

the complex prior to eruption or emplacement at higher levels. The trend to higher K2O with

increasing SiO2 may be interpreted as an anhydrous fractionation trend resulting from the continued

loss of volatiles during volcanic eruptions.

The second trend which is defined by the intrusions related to mineralization, is marked by a greater

enrichment of silica and is best explained by fractionation of plagioclase and biotite (Figure B.3).

Figure B.3

39

This trend is considered to mirror a higher partial pressure of water in the chamber at the time the

porphyry deposits formed and suggests a key element in the sequence of events required to produce a

porphyry deposit is the containment and controlled release of the volatiles.

0

2

4

6

8

K 2O

(w

t %)

A

High pressure fractionation( pyroxene, plagioclase, olivine)

Low pressure fractionation(olivine, plagioclase, pyroxene)

SiO2 (wt %)45 50 55 60 65 70 75

Mineralizing trend(biotite, plagioclase fractionation)

Figure B.4

40

Part C

Hydrothermal Systems in Sedimentary Basins andMetamorphic Terranes

C1: Key Issues

Development and maintenance of fluid reservoirs.

Extent to which fluids are sourced within basins or within basement.

Mechanisms of fluid release.

Mixing processes.

Low temperature transport of reduced sulfur and metals.

C2: Low Temperature Hydrothermal Systems

The discussion of the genesis of low temperature (sub 200oC) metal deposits pivots on two relatedquestions:

1. The solubility of base-metal sulfides in hydrothermal fluids at temperatures below

200 to 250oC.

2. The rate at which sulfate may be reduced to sulfide in hydrothermal solution below

200 to 250oC.

Experimental studies on the solubility of sphalerite and galena have shown that it is very difficult to

dissolve sufficient metal and sulfide in reduced solutions at temperatures below about 200oC and

therefore difficult, if not impossible, to transport sufficient metal and sulfide in the one fluid. Yet

many Pb-Zn deposits are known to form at low temperatures.

In contrast the base metals, Pb, Zn and Cu are readily soluble in sulfate-rich solutions at low

temperature. However, because the rate of reduction of sulfate to sulfide is very slow at

temperatures below 200 to 250oC (see the summary of experimental studies by Ohmoto and Lasaga,1982) it is difficult to form abiogenic sulfides from sulfate-rich solutions.

41

These fundamental difficulties has never been adequately resolved. The family of proposed

solutions (see Anderson and McQueen; 1982) include:1. Transport of metals in sulfate-rich brines and deposition by mixing with H2S-rich fluids.

2. Transport of metals and H2S in an acidic fluid at high partial pressures of CO2 to maintain

dolomite saturation.

3. Transport of metals in sulfate-rich brines and deposition by in situ thermo-chemical reduction of

sulfate to sulfide.

Many workers have argued that sulfide may be produced from sulfate through thermo-chemical

reduction in sedimentary basins at low temperatures. The strong argument for this is the

coexistence of sulfides and sulfates where the sulfide mimics the sulfur isotope ratios of the sulfate.

The weakness of this mechanism is the slow rate of reduction of sulfate to sulfide at low

temperatures.

Temperature, 0C600 300 200 100 50 25 0

(SO42--H2S)

0 /00

40

30

20

10

0

Eq. curve

Up. Miss. Valley

SE MissouriFinlandiaPark Ci ty

Porphyry Cu

Creede

Kuroko

Carl in

Hansonburg

Figure C.1

42

Data on the partitioning of sulfur isotopes between sulfides and sulfates in natural systems also

suggests sulfate and sulfide species do not equilibrate at low temperature. In high temperature

systems such as porphyry Cu-Au deposits, isotopic equilibrium between sulfate and sulfide is

established (see Figure C.1) but in low temperatures systems, such as Mississippi Valley deposits,equilibrium is not established. Such data argue against in situ reduction of sulfate (point 3) as adominant mechanism for precipitating sulfides.

The easiest explanation of low temperature sulfide deposits is that the metals and the reduced sulfur

were transported in two separate fluids which mixed at the site of ore formation. Recent studies on

the Pb-Zn deposits of south-east Missouri (Goldhaber et al., 1995; see below) make a strong casefor the mixing mechanism.

Pyrite as an indicator of non-equilibrium processes

The distribution of pyrite may provide some information about approach to equilibrium in low

temperature systems. In particular, the presence or absence of pyrite may provide some clues about

temperatures of the system and about loss of gas pressure.

Pyrite is unique among the common sulfides in that the sulfur in pyrite is partially oxidized. Its

precipitation depends on the presence of S2(2-)

Fe2+ + S2(2-) FeS2

S2(2-) is an intermediate species in the stepwise reduction of sulfate to sulfide.

H2S S2(2-) S2O3(2-) SO4(2-)

If sulfate cannot be readily reduced, the precipitation of pyrite is inhibited. This is most likely to

occur in the temperature window between about 80 and 250oC (see Figure C.2).

At temperatures below about 80, bacterial reduction can produce pyrite, for example, py1 and py 2

at McArthur River. At temperatures above about 250oC, pyrite is readily precipitated through

43

thermo-chemical reduction processes. Pyrite is a common phase in most intermediate temperature

systems (around 250 - 400oC).

4

0

-4

-8

100 200 300 400

Non-equil ibr ium Equil ibr ium

Biogenic PyritePyrite precipitat ioninhibited

RedoxScale

Redox Scale = log [ total sulfate/total sulf ide]

Temperature oC

Hemat ite

Pyrrhot ite

PyriteMagnetite

To precipitate pyrite in the temperature interval between 80~250oC there needs to be a specific

reaction path. Hence, a general absence of pyrite at the stage of base metal sulfide deposition in

sediment hosted deposits hints at low temperature systems, most probably sub 200oC.

Vapour phase separation in a reduced fluid may promote pyrite precipitation by the reaction:

2H2S S2(2-) + 2H+ + H2(g)

Hence the distribution of pyrite in deposits which form generally at temperatures below 200oC

could be used to indicate domains/periods of pressure release.

Figure C.2

44

Reef

Bonneter re Fm

Precambr ian

Lamotte Fm

Dav is Shale

0 25km

N

OLD LEAD BELT

VIBURNUM TREND

Casteel Mine

Mining districts

Structures ut i lized by sour gasfrom basement

Figure C.3

45

C3: Mississippi Valley Pb-Zn deposits - A useful analogue for lowtemperature systems

The Mississippi Valley deposits of the mid-continent United States are perhaps the most intensely

studied low temperature hydrothermal deposits. Of particular significance is the progress made in

understanding the nature of the fluid reservoirs in the Pb district of Southeast Missouri. This district

represents the worlds largest Pb resource.

Some salient geological features of the S.E. Missouri Pb district:

Deposits are located within the reef complex of a Cambrian carbonate sequence, the Bonneterre

Formation, which onlaps a Precambrian massif; the St. Francois Mountains. The major districts(Old Lead Belt, Viburnum Trend) are distributed around the massif (see Figure C.3).

The basement is covered by the Lamotte Sandstone and the deposits overlie a pinchout in this

unit and are overlain by the Davis Formation (see inset to Figure C.3).

E-W trending structures across Kentucky, Illinois and Missouri and eastern Kansas (part of the38 parallel lineament of Heyl, 1983) represent a basement weakness. This fracture system cutsacross the northern part of the St. Francois Mountains.

The sulfide mineralogy in Mississippi Valley Type deposits is dominated by galena and

sphalerite. The Casteel Mine in the northern Viburnum Trend is an example of a chalcopyrite-

rich deposit (Dunn, Jr., and Grundmann, Jr., 1989).

A section through the Casteel Mine is shown in Figure C.4. Chalcopyrite-rich ore (averaging5.5%Cu) with minor galena is overlain by pyrite(marcasite)-, galena-, chalcopyrite-rich ore. Thepyrite-rich ore tends to occur at the contact with the Davis Formation and commonly passes

abruptly into the underlying chalcopyrite-rich ore. Quartz is a common gangue mineral in theCasteel Mine and minor bornite occurs.

46

Reef

D a v i sF o r m a t i o n

B o n n e t e r r eF o r m a t i o n

P r e c a m b r i a nB a s e m e n t

50m

C r o s s - S e c t i o n C a s t e e l M i n e

Pyrite - (marcasite) -galena - chalcopyriteOre

Chalcopyrite Ore

Grainstones +shaley mudstones

Conglomerate

Domain of greatest pressure loss

Domainof

initial mixing

Some salient geochemical features of the S.E. Missouri Pb district:

Temperatures of ore formation were around 100-150oC and brine salinities around 15-30 weight

percent.

Pb isotope data (Figure C.5; Goldhaber et al., 1995 show that Pb in the main stage ofmineralization (cube-octahedral galena) in the Viburnum Trend and the Old Lead Belt has a lowratio of 208Pb/204Pb for a given 206Pb/204Pb compared to other mineral districts (Illinois -Kentucky Fluospar District, Upper Mississippi Valley, Northern Arkansas, Tri-State and Central

Missouri) and to late stage mineralization (cubic galena) in the Viburnum Trend.

The regional survey of Goldhaber et al. (1995) has shown that the Pb isotope signatures of theIllinois - Kentucky Fluospar District, Upper Mississippi Valley, Northern Arkansas, Tri-State

Figure C.4

47

N

Ouachita Mountains

43

42

41

40

39

38

18 19 20 21 22 23 24

206Pb / 204Pb

208 P

b / 2

04Pb

Upper Mississippi Valley

Il l inois/KentuckyFlourspar distr ict

Central Missouri,Tri-state

&Northern Arkansas

octahedral and disseminated galena cubic galena

Viburnum Trend

Il l inoisBasin

Arkoma BasinReelfoot

Rift

Forest CityBasin

Old Lead beltCentral Missouri

Viburnum Trend

Tri-State

0 200 km

Figure C.5

Figure C.6

48

and Central Missouri and of the late stage mineralization (cubic galena) in the Viburnum Trendare consistent with basinal sources for the Pb (Illinios, Arkoma Basins). In contrast the Pb ofthe main stage of mineralization in the Viburnum Trend and the Old Lead Belt appears to have

been derived from the Lamotte Sandstone and/or underlying weathered basement.

There is a broad correlation between the S and Pb isotope data (Sverjensky et al., 1979;Goldhaber et al., 1995). The isotopically heavy sulfur of the Viburnum Trend and the Old LeadBelt correlate with less radiogenic values of 206Pb/204Pb (compare Figures C.5 and C.7; S34> +10). Isotopically lighter sulfur (-10 to +10) correlates with the more radiogenic values of206Pb/204Pb (Illinois - Kentucky Fluospar District, Upper Mississippi Valley, NorthernArkansas, Tri-State and Central Missouri).

500

300

100

-10 0 10 20

34 S

Old Lead Belt

Viburnum Trend cubic oct

Central Missouri

Northern Arkansas

Tri-State

Metresabovebasalsandstone

Given the correlation between the Pb and S isotopes, it appears that the isotopically lighter

sulfur (-10 to +10) was sourced from the basins and the isotopically heavy sulfur ( S34 >+10) sourced from the basement (see also data of Goldhaber et al.,1995). In the area aroundthe St. Francois Mountains isotopically heavy sulfur occurs within the upper Bonneterre

Formation below the Davis Formation.

Figure C.7

49

The main stage of mineralization (octahedral galena) in the Viburnum Trend is characterised bybrines with high values of K/Cl (Figure C.8) as well as high ratios of Mg/Cl and Br/Cl (Viets etal., 1989).

0.07

0.05

0.03

0.01

0 80 160 240 320

Kilometres

K/Cl

Northern Arkansas

Seawater Tri-State

ViburnumTrend

cubic

octahedral

CentralMissouri

The Genesis of the Pb-Zn Deposits of S. E. Missouri

Deposits of the Viburnum Trend and the Old Lead Belt were formed from mixing fluids from

two reservoirs that developed within the basement (Figure C.9). This contrasts with other (lesssignificant?) mineral districts in the mid-continent that were developed from basinal reservoirs.

The Lamotte Sandstone was a major flow path for Pb-Cu-Zn-rich but H2S-poor fluids from areservoir within the weathered basement the Lamotte Sandstone. This fluid most likely

contained metastable sulfate.

A H2S-CH4(?)-rich gas (sour gas) mixed with the brine from the Lamotte aquifer to causesulfide deposition. It is suggested that this gas was also tapped from a basement reservoir. The

heavy sulfur isotope signature of the main stage mineralization suggests that the gas was derived

Figure C.8

50

from a reservoir in which basinal brines underwent extensive reduction at temperatures above

200-250oC.

H2S vapor from reservoir ofhighly reduced basinal brines

Pb-Zn rich fluids-part ial ly reduced

Basinal oil &gas- light sulfur isotopes

- source of heavy sulfur isotope signature

Ore deposit ion- above pinch out

basement faults

Aquiclude -Davis Shale

Heavy sulfur isotopesignature in dispursed sulf idesbeneath aquiclude

It is suggested that the underpinning regional control on the south-east Missouri Pb district was

the interaction between E-W trending basement structures and the St Francois Precambrian

massif during deformation, probably in the Late Devonian or Early Carboniferous. Basement

fluids were focused into domains of low pressure around the massif; the metal-rich brines

utilized the Lamotte aquifer and the sour-gas utilized steeper structures which directly penetrated

the basement.

Figure C.9

51

The Davis Formation acted as an aquiclude. Ore deposition occurred where the Davis Formation

was ruptured above the pinchouts with the Lamotte Sandstone. The chalcopyrite-rich ores

represent the point of initial mixing of brine and gas. The pyritic ores represent domains of

substantial vapor loss at the contact with the Davis Formation.

Remarks: A key part of the ore-forming process was maintenance of the integrity of the metal-rich, H2S-

poor reservoir(s). Inmixing of H2S, into this reservoir(s) led firstly to a loss of Cu and secondlyPb from the fluids.

The heavy sulfur isotope signature (> 10; Goldhaber et al., 1995) within the upper BonneterreFormation, just below the Davis Formation, may reflect partial degradation of the metal-richfluid reservoir through leakage of sour gas into the reservoir. The gas would have been

concentrated beneath the aquiclude leading to sulfide enrichment in this horizon (Coveney,1989). This suggests a greater potential for low temperature Cu deposits where the integrity ofthe reservoirs are maintained.

The regional Pb isotope studies of SE Missouri provided the strongest evidence for the sources of

the ore fluids. Correlating these data with the S isotope data and cation ratios yields some

empirical arguments which may be useful in the absence of definitive Pb isotope data.

Consistently heavy sulfur isotopes values and high values of K/Na, K/Cl and Br/Cl may be

considered indicative of well-mixed basement sources.

There is some indication that heavy sulfur isotopes values may be correlated with mineralizing

systems related to major structures.

52

C4: Some constraints on the Proterozoic Pb-Zn deposits of NorthernAustralia

These deposits were most likely formed at similar temperatures (100-1500C) to the MVTdeposits.

The lack of pyrite at the base-metal stage is an important clue to their low-temperature origins.

Sulfide precipitation most likely occurred by mixing of base-metal rich and H2S rich fluids; by

analogy with the south-east Missouri Pb district.

In contrast to the MVT deposits the individual deposits are very large. This property suggests

steep structures controlled the fluid flow of both the metal-rich and the H2S-rich fluids.

Also in comparison with the south-east Missouri Pb district, both the heavy and light sulfur

isotope ratios are missing from the main stages of mineralization.

This suggests that specific reservoirs which could generate extreme sulfur isotope compositions

were not important; either within the basin or basement.

The sulfur supply was well mixed and may reflect basement as well as basin sources.

basemetals

McArthurRiver

H2S-richgas

Figure C.10

53

C5: The Mount Isa Cu Deposit

The Mount Isa Cu orebodies are high-grade, chalcopyrite-rich and average 3% Cu with massive

tabular zones of up to 25% Cu. Most recent workers (Heinrich et al., 1995; Heinrich et al., 1989;Andrew et al., 1989; Perkins, 1984; Swager, 1985) consider that these orebodies formed at a latestage in a regional tectonic and metamorphic event, the Isan Orogeny. However, the processes

which led to the development of the Cu orebodies, which to a first approximation might be

considered a giant mass of chalcopyrite, and their relationship with the adjacent Pb-Zn ores remaincontroversial. How does nature form a gaint mass of chalcopyrite? In this discussion some parallels

will be drawn with the chalcopyrite-rich parts of deposits of the Viburnum Trend and the MVT

model discussed above will be utilized in developing a low-T hydrothermal model for the Mount Isa

Cu deposits.

Eastern CreekBasalt

NS

SilicifiedZone

4600mN

6400mN

Perkins (1984)

Some salient features of the geology of the Mount Isa Cu deposit The Cu orebodies are located within a sequence of steeply dipping, N-S trending dolomitic

shales. They lie above a faulted contact with the Eastern Creek Volcanics which are extensively

altered (Figure C11).

Figure C.11

54

It is considered that the overall direction of fluid flow within the Cu ore-system was from south

to north and upwards from the contact with the Eastern Creek Volcanics, the Buck Quartz Fault(Figures C.11 and C.12; Perkins, 1984; Waring, 1990). The overall zonation in oxygen isotopesalong the focus of fluid flow (Waring,1990) is from light values in the south and at depth (O18

around 10) to heavier values in the north and higher in the section (O18 around 13 ).

2500

4548

5553

5033

1520

14

15

16

13

17

Mount Isa Fault

Posit ion of 1100 onBuck Quartz Fault

4 Km

5553

5033

4548

2500

1520

Oxygen Isotope Contours (per mil)

After Waring (1990)Surface Plan Long Sect ion

Eastern Creek Volcanics

Figure C.12

55

In any one section the pattern is from lighter values at the focus of fluid flow to heavier values in the

wall rocks (see the plan view). This pattern is consistent with progressive reaction of fluids androck.

3270

3880

5350

4950

2500

2000

4550

S

N1100 CU OREBODY

ALTERATION

3D PROJECTION

PYRITE

DOLOMITICALTERATION

SILICEOUSALTERATION

TALCWaring (1990) Figure C.13

56

The distribution of mineral assemblages within the Cu ore-system in part reflects a S to N

zonation and in part temporal variation. Domains of Mg-talc and chlorite alteration are best

developed to the south of the orebodies (Figure C.13; Waring, 1990). Domains of dolomiticalteration are best developed subjacent to and above the Cu orebodies. Siliceous alteration isfocused on the Cu orebodies and postdates the dolomitic alteration.

Pyrrhotite occurs as a marginal phase to the Cu orebodies (Perkins, unpublished manuscript) andthere are more limited domains of coarse grained pyrite on the periphery of the Cu deposits.

Greenstones

Total dolomitization

3400m RL

3000m RL

2600m RL

Coarse pyr i te

Pyrrhoti te

Buff a lterat ion and bleaching

Domains of greatest pressureloss

Diagram af ter Perkins (unpubl ished manuscript )

5000

m N

7400

m N

6600

m N

5800

m N

Distinct assemblages of biotite-pyrrhotite-stilpnomelane, K-feldspar, ferroan dolomite and

sulfides occur to the north of the Cu orebodies and between the Pb-Zn lodes. Eastern Creek

Volcanics below the mine are highly altered to an assemblage of Mg-chlorite and rutile.

Figure C.14

57

Some salient features of the geochemistry of the Mount Isa Cu deposit

The fluid inclusion study of Heinrich et al. (1989) yielded homogenization temperatures

between about 150 and 250oC for dolomite and 125 and 200oC for quartz associated with

chalcopyrite (Figure C.15). There is also a change from Ca-rich to Na-rich fluids from thedolomitic to siliceous stages of alteration. The least deformed quartz veins associated with Mg-

chlorite assemblages below the Mine yielded homogenization temperatures mostly between 200

and 250oC and Na-rich fluids.

50 100 150 200 250 300Temperature of homogenization 0C

30252015105

105

Mg chlorite-ruti le alterat ion

Quartz chalcopyrite ore

Silica dolomite alterat ion

Dolomit ic alteration

CaCl2-rich inclusions (Tice < -230C)NaCl-r ich inc lusions (Tice > -230C)

Numberofobservat ions

S34 values range from around +10 to +25 with consistently heavy values occurring in thesouthern lodes. Consistently heavy values (>+20) also occur in the southern extremities ofthe Cu ore-system and in the Mg-chlorite altered Eastern Creek Volcanics beneath the mine

(Figure C.16; Andrew et al., 1989; Heinrich et al.,1995).

Figure C.15

58

10

5

0

20

10

0

5

0-20 -10 0 10 20

34 S (permil CDT)

Number o foccurr

A Metabasalts east of Mount Isa

B Mount Isa copper ores

C Metabasalts below mine

Least altered metabasaltEpidote-sphene alterationCarbonate-Fe oxide alterat ion

Minor sulfide south of mineMajor orebodies

Moderately al teredMg chlorite-ruti le alterat ion

C13 values for secondary dolomite range between about -3 and -10 and there is a broadpositive correlation with O18 (Figure C.17; Waring, 1990). The most extreme low values ofO18 and C13 are also returned from minor carbonate in the Mg-chlorite alteration below themine (Heinrich et al., 1995).

Br/Cl ratios of saline fluid inclusions are high (Figure C.18; Heinrich et al., 1995) andcomparable to values for crystalline basement brines. The main stage of mineralization in the

Viburnum Trend is also characterised by brines with high values of Br/Cl.

Figure C.16

59

13C

18O

"Orebody" Dolomite Isotopic Composit ion

Sth Nth

Orebody SouthOrebody 4200 mNOrebody 5030 mNOrebody Nor th (650)

-1

-3

-5

-7

-9

-119 13 15 17 19 21 23

Development of a low-T genesis model for the Mt Isa Cu ores

Origins of the mineral zonation in the Cu ore-system The zonation in mineral assemblages in the Cu ore-system may be interpreted in terms of a

thermally declining hydrothermal system from S to N and with time. From the log mSiO2 vs

temperature diagram (Figure C.19) the highest temperature assemblage would be the Mg-chlorite -rutile assemblage in the Eastern Creek Volcanics below the mine followed by the Mg-talc - chlorite

assemblages in the southern part of the system (the positions of the phase boundaries in the diagramwill vary according to the values of the solution parameters noted. The topology will be invariant).As the system cooled from S to N and with time the dolomitic alteration developed. Further cooling

lead to the siliceous alteration and formation of the Cu orebodies. In this interpretation the size of

the hydrothermal system contracted with cooling and time.

Figure C.17

60

Sedimentary hal i te

Basin brinesBulk hal i te di ssolution brines

Residual b i tterns

0.0001 0.001 0.01 0.1

Bromine/chlorine (wt ra t io)

SMOW

Magmati c f lu ids

Crystal l ine basement bri nes

Mount Isa

Mine

Metabasal ts

Quartz-chalcopyrite ore

Epidote-sphene al teration

Carbonate-Fe oxide altn

Figure C.18

61

-3.5

-3.0

-2.5

-2.0

100 200 300

Temperature (deg.C)

Dolomite

Talc

Clinochlore

Mg-talc zones south of the deposit

Mg-chlorite- rutile

alterationin

basaltbelow

Mt Isa Mine

QuartzSaturation

log aAl3+/(aH+)3 = -2.15 log aCa2+/(aH+)2 + 2 log CO2 = 5.3

DolomiteAlteration

Quartz &chalcopyrite

Temperature of chalcopyrite deposition and implications Genesis models for the Mount Isa Cu ores pivot on the question of the temperature of chalcopyrite

deposition. The raw homogenization temperatures indicate quartz and chalcopyrite deposition at

low temperatures - below 200oC and higher temperatures of deposition for the dolomite. Should a

pressure correction be applied to these data to obtain higher temperatures for chalcopyrite

deposition; the procedure adopted by Heinrich et al. (1989)? The presence of a coexisting gasphase would argue against applying a pressure correction and rare methane-rich inclusions have

been observed.

Figure C.19

62

The above explanation of the mineral zonation also implies a decline in temperature from the stage

of dolomite deposition to the stage of quartz and chalcopyrite precipitation. There is some

consistency here between the fluid inclusion and mineral zonation data. In the following discussion

the raw homogenization temperatures are taken as a reasonable approximation to the true

temperatures of chalcopyrite deposition. This has some significant implications and parallels may

be drawn with low-T chalcopyrite-rich deposits of the MVT class and analogous models of

formation developed. The raw homogenization temperatures are probably minimum estimates

of the temperature of chalcopyrite deposition. The very few samples of quartz-chalcopyrite studied

Mt Isa Cu

1010102103 1 km

10

km

1

100 Mantle

CH4 plus H2S gas der ivedf rom deep-crusta l reservoir - sulfur f r om completereduct ion o f sulpha te sul fur

Eastern CreekVolcanics

Basement reservo i r o fevo lved bas ina l br ines- less 200 oC- metas tab le su l fa te- H2S poor- base meta ls

Figure C.20

63

by Heinrich et al. (1989) that contained methane-rich vapour inclusions also gave high

homogenization temperatures - up to 240 oC - which may be more indicative of the true

temperatures of chalcopyrite deposition. The arguments presented below would be valid up to

temperatures of 250 to 275 oC.

A low-T mixing model

It is possible to interpret the geological and geochemical data in terms of a two fluids mixing at low

temperature (Figure C.20).

One of these fluids was a metal-rich, H2S-poor, saline brine, possibly containing metastable

sulfate. It was sourced from a low-T reservoir at a relatively high level in the crust, its Cu- rich

nature reflecting reaction with the Eastern Creek Volcanics.

The other fluid was a CH4 - H2S - rich gas (at least at the T and P conditions which prevailed atthe site of ore formation) that was sourced from high-T reservoir much deeper in the crust. Inthis reservoir basinal brines were extensively reduced to yield a heavy sulfur isotope signature.

The low values of C13 in secondary carbonate are indicative of a CH4 - rich fluid.

The assemblage of biotite, pyrrhotite, stilpnomelane, K-feldspar, ferroan dolomite to the north of

the Cu orebodies and between the Pb-Zn lodes is also an indication of a reduced and alkaline

fluid. This assemblage predates the main stage of mineralization (Perkins, unpublishedmanuscript).

The mixing of brine and CH4 - H2S - rich gas could cause the precipitation of chalcopyrite while

simultaneously dissolving preexisting dolomite and sulfides.

The hydrolysis of CH4 is a potential source of acidity to drive dissolution reactions:

CH4 + 3H2O HCO3-

+ 3H+ + 4H2

The precipitation/dissolution of chalcopyrite will be controlled by the relative changes to H2S, CH4and H+:

64

8Cu+ + 8Fe

2+ + 16H2S + HCO3

-

8CuFeS2 + CH4 + 23H+ + 3H2O

The addition of H2S will be very effective providing the pH is buffered by dolomite dissolution.

Regional-scale flow paths It is suggested that the flows path for the saline brine which carried the copper were the west-

dipping structures seen in the reflection seismic profile of the Mount Isa Inlier (Figure 7 in Golebyet al., 1996). The point of entry marked by the Mg-chlorite - talc zones of alteration.

The input of the CH4 - H2S - rich gas input was possibly focused to the north of the Cu orebodies

and marked by the development of the distinct assemblages of biotite-pyrrhotite-stilpnomelane, K-

feldspar, ferroan dolomite and sulfides.

The gas is considered to have evolved from deep-seated crustal fluids potentially rich in Pb, Zn, Cu

and Au. The flow paths into the upper crust were the steeply east-dipping structures in the

reflection seismic profile of the Mount Isa Inlier (Goleby et al., 1996; Figure 7).

Site of ore formation - disruption of a regional seal?

The site of ore deposition is considered a site at which a regional pressure seal was disrupted.

Fluids (brine and gas) were consequently focused into this site. Following the MVT analogy andargument about the significance of pyrite in low-T systems, it is suggested the domains of coarse

grained pyrite documented by Perkins (Figure C.14) may be considered the sites of lowest pressureand substantial vapor loss.

65

Part D

Crustal Scale Systems, the Third Dimension and Lode Gold Deposits

Moho

Impermeablelayer

H2O, O2, CO2, NaCl, SO42-

CH4, CO2, H2S

Reservoirs of CH4 - H2S f luids,Coexist ing with saline fluidsRich in Au - Pb - Zn - Fe

Reservoirs ofCO2-H2O fluids

Reservoirs ofsaline fluid &CO2-rich fluids

Figure D.1

66

D1: Key issues

Scale at which fluids are moving within the crust and depth of hydrothermal systems.

Compositions of fluids with depth, temperature and composition of host rocks.

NaCl - CO2 - CH4 - H2O phase relations at high temperature and pressure.

Transport and deposition of metals at high temperature and pressure.

Development of homogenous fluid reservoirs.

Mechanisms of fluid release.

D2: The third dimension

Major hydrothermal ore deposits are commonly associated with faults or lineaments that maybe traced 10s to 1000s of kilometres on the earths surface. Dimensional argument suggests

that the third dimension of hydrothermal systems centered on such structures could be of the

order of 10s km. Recent deep seismic reflection profiles of the Mount Isa Inlier (Goleby etal., 1996; Goncharov et al., 1996) and the Eastern Gold Fields (Drummond and Goleby,1993) have imaged seismic structures of 5 - 50 km length within the crust with somestructures penetrating the mantle. These profiles may be providing an image of the third

dimension of hydrothermal systems.

How will fluids be distributed in the third dimension, what will be their chemical character

and what will be their ore-forming potential? Fluids may be contained within discrete

reservoirs separated by regional-scale seals. Recent theoretical argument (Hobbs and Ord,1997) suggests that convection within reservoirs will be an important mechanism forgenerating large volumes of fluid of constant composition. Mixing of fluids from reservoirs

of contrasting fluid composition is likely to be a significant mechanism for generating large

tonnage, high grade mineral deposits.

There will be some systematic variations in compositions of reservoir fluids with depth,

temperature, nature of host rocks and tectonic setting. Salt-rich fluids will occur in reservoirs

within or below sedimentary basin or in reservoirs that develop within and around magma

chambers. Sulfate is likely to dominate over H2S in low - T reservoirs (less than about

67

200C). Reservoirs of CO2 - H2O fluids will commonly occur in the mid to lower crust and insome magmatic systems.

The lower crust may well contain reservoirs of water-poor fluids and some of these may be

dominated by CH4 rather than CO2 with H2, N2 and H2S. CH4-N2 - rich fluids in the mid to

lower crust may well coexist with a saline brine. Such fluids may develop at times of re-

hydration of the mid to lower crust by reduction of CO2-bearing surface waters. In a low

fluid/rock environment in the lower crust re-hydration of the crust could lead to H2O depleted

fluids that are enriched in other volatile components. Loss of H2O would lead to CH4enrichment relative to CO2 according to the reaction

H2 + CO2 CH4 + 2H2O

Dahomeyan

belt type granitoids

basin type granitoids

sedimentary basins

Togo series

Tarkwaian

volcanic beltsBirimian

ASHANTI BELT

60 Km

Aya nfur i

Bogosu

Ashanti

Konongo

Prestea

Figure D.2

68

Reservoirs of water-poor fluids could develop during periods of extension following periods of

compression, heating and crustal dehydration.

D3: The Ashanti Au Belt, Ghana, West Africa Windows into the fluids of the deep crust

The Ashanti Gold Belt, a 250km long, NE trending, structural belt that divides marine sedimentary

rocks and volcanic greenstone belts of the Palaeoproterozoic Birimian Supergroup (Figure D.2).The deposits within the belt are Early Proterozoic examples of lode Au