Sequence Stratigraphy of Lacustrine

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SEQUENCE STRATIGRAPHY OF FLUVIAL AND LACUSTRINE DEPOSITS INTHE LOWER PART OF THE CHINLE FORMATION, SOUTH CENTRAL UTAH,

UNITED STATES: PALEOCLIMATIC AND TECTONIC IMPLICATIONS

A THESISSUBMITTED TO THE FACULTY OF THE GRADUATE SCHOOL

OF THE UNIVERSITY OF MINNESOTABY

Joseph John Beer

IN PARTIAL FULFILLMENT OF THE REQUIREMENTS

FOR THE DEGREE OFMASTER OF SCIENCE

December 2005


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Copyright Joseph John Beer 2005


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ABSTRACT

Field and well log data from central Utah identify three sequence bounding

unconformities in the lower portion of the Chinle (Shinarump, Monitor Butte, Temple

Mountain, and Moss Back Members) which function to divide these deposits into

three periods of incision and subsequent valley fill. The initial period of degradation

is marked by interfluve paleosols and truncation of the underlying Moenkopi

Formation creating a paleovalley which constrains the deposition of these four

members. The first paleovalley fill, represented by the Shinarump Member, is

interpreted as a confined sandy low-sinuosity river system. A second period of

incision is marked by truncation of the Shinarump and correlative paleosols and

pedogenically modified strata. The Monitor Butte and correlative Temple Mountain

Members overlie this unconformity and consist of mudstones and sandstones

representing fluvio-lacustrine deposition and vertisols and interbedded mudstones and

sandstones deposited in a high-sinuosity river system. A final cut-and-fill cycle within

the pre-Shinarump paleovalley is filled by the high- and low-sinuosity fluvial deposits

of the Moss Back Member. Truncation of the Monitor Butte and Temple Mountain

Members and the Moenkopi Fm. and interfluve pedogenesis mark the preceding

surface of degradation.

Erosional unconformities and correlative extensive pedogenesis within the

Chinle in central Utah indicate a depositional history involving alternating periods of

landscape degradation and aggradation. The tectonic setting of the Chinle basin

within a dynamically subsiding back-arc basin may provide a mechanism for 1) uplift

and erosion of Lower Triassic strata creating the master paleovalley surface, and 2)

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later subsidence of the Chinle depositional basin. The high frequency, low amplitude

cut-and-fill nature of the lower portion of the Chinle is likely the result of changes in

sediment flux and discharge as evidenced by detailed facies and pedogenic analysis,

and paleoecological and paleoclimatic data. In addition, the nature of deposition

within an incised valley network had a large effect on accommodation, and the spatial

distribution of facies preserved in the lower portion of the Chinle Formation.

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AKNOWLEDGEMENTS

First and foremost I wish to thank Tim Demko for the first-hand introduction

to the geology of the Colorado Plateau and for instilling within me the desire to visit

every Mesozoic outcrop preserved on the Colorado Plateau. Tims personality as well

as his interest and knowledge pertaining to the stratigraphy of the Plateau made a lack

of discussion, feedback, and encouragement an impossibility.

I also owe a large debt to both Corey Wendlund and Ryan Erickson for the

many weeks each of them spent assisting me with fieldwork. I would like to thank

Marsha Meinders-Patalke, Erik Gulbranson, and Riyad Ali-Adeeb for numerous

thoughtful discussions throughout the course of this project. I also wish to thank John

Swenson, and Julie Etterson for thoughtful comments and insights during the editing

process.

This project was funded in part thanks to the generosity of the Colorado

Scientific Society, the University of Minnesota Duluth Department of Geological

Sciences, AAPG Grants in Aid, and the University of Minnesota VDIL. Digital well

logs were provided by an educational grant from MJ Systems.

Lastly I would like to thank my wife Jenny and the rest of my friends and

family for their patience, support, and encouragement throughout the course of this

project.

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TABLE OF CONTENTS

SIGNATURE PAGE

TITLE PAGE

COPYRIGHT NOTICE PAGE

ABSTRACT i

ACKNOWLEDGEMENTS iii

TABLE OF CONTENTS iv

LIST OF FIGURES vii

CHAPTER I: INTRODUCTION 1

1.1 Purpose 1

1.2 Location.. 2

1.3 Methods... 5

1.3.1 Field Methods.. 5

1.3.2 Petrographic Methods 5

1.3.3 Subsurface Methods. 6

CHAPTER II: GEOLOGIC SETTING. 7

2.1 Tectonics.... 7

2.2 Stratigraphy.... 11

2.2.1 Pennsylvanian Permian. 12

2.2.2 Early Triassic 12

2.2.3 Late Triassic. 14

2.2.4 Jurassic... 22

2.3 Paleogeography.. 22

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2.4 Paleoclimate..... 25

2.5 Paleoecology... 29

CHAPTER III: DATA AND RESULTS 30

3.1 Unit Descriptions Lithology, Facies, and Architecture 30

3.1.1 Basal Unconformity... 30

3.1.2 Shinarump Member... 32

3.1.3 Post-Shinarump Unconformity. 33

3.1.4 Monitor Butte and Temple Mountain Members 34

3.1.5 Pre-Moss Back Unconformity....... 39

3.1.6 Moss Back Member... 41

3.1.7 Petrified Forest Member........................... 45

3.2 Paleocurrent Data... 47

3.3 Petrographic Data... 49

CHAPTER IV: SEQUENCE STRATIGRAPHY... 51

4.1 Introduction to Sequence Stratigraphy 51

4.2 Application to Continental Strata 52

4.3 Discussion of Unconformity Paleosols 55

4.4 Sequence Stratigraphic Interpretation. 57

4.4.1 Sequence Boundary 1 60

4.4.2 Depositional Sequence 1 60




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CHAPTER V: CONCLUSIONS.... 64

5.1 Depositional History... 64

5.1.1 Landscape Degradation 1 64

5.1.2 Landscape Aggradation 1.. 66

5.1.3 Landscape Degradation 2... 66


5.1.5 Landscape Degradation 3.. 68


5.2 Lithostratigraphic Implications... 70

5.2.1 Correlation of the Temple Mtn. and Monitor Butte Members.. 70

5.2.2 Suggested Modifications to Lithomember Boundaries. 70

5.3 Summary................. 76

REFERENCES CITED 77

APPENDIX I: MEASURED SECTION DATA... 86

APPENDIX II: PALEOCURRENT DATA.. 153

APPENDIX III: PETROGRAPHIC DATA... 169

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LIST OF FIGURES

Figure 1.2-1 Distribution of Triassic outcrops on the CO Plateau 3

Figure 1.2-2 Map of the field area. 4

Figure 2.1-1 Map of the tectonic setting during the Late Triassic 9

Figure 2.2.2-1 Stratigraphy and nomenclature of the field area 11

Figure 2.2.3-1 Triassic timescale.. 16

Figure 2.2.3-2 Map depicting location of Painted Desert Paleovalley.. 18

Figure 2.3-1 Map showing location of field area during the Late Triassic... 23

Figure 2.3-2 Paleogeographic reconstructions of Late Triassic 25

Figure 3.1.1-1 Unconformity paleosol and lacustrine clinoforms near MS 27. 31

Figure 3.1.4-1 Lacustrine turbidite facies. 37

Figure 3.1.4-2 Rooted horizon in Monitor Butte Member 37

Figure 3.1.4-3 High sinuosity fluvial facies in Monitor Butte Member... 38

Figure 3.1.5-1 Pedogenic carbonate nodules from unconformity paleosol... 40

Figure 3.1.5-2 Intense bioturbation in unconformity paleosol.. 41

Figure 3.1.6-1 Aerial photograph of a portion of Blue Notch Canyon. 43

Figure 3.1.6-2 Mon. Butte, Moss Back, and Pet. Forest Mbrs. At MS 16 44

Figure 3.1.7-1 Trough cross-beds in Petrified Forest Member. 46

Figure 3.1.7-2 High sinuosity fluvial deposits in Petrified Forest Mbr 47

Figure 3.2-1 Maps depicting paleocurrent measurements 48

Figure 3.3-1 QFL diagrams of sandstones from the 3 depositional sequences. 50

Figure 4.1-1 Glossary of sequence stratigraphic terminology.. 51

Figure 4.2-1 Model depicting controls upon fluvial incision vs. aggradation.. 53

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Figure 4.2-2 Lacustrine classification scheme.. 54

Figure 4.4-1 Stratigraphic cross section of the lower portion of the Chinle Fm... 58

Figure 4.4-2 First order sequence stratigraphic framework... 59

Figure 4.4-3 Second order sequence stratigraphic framework.. 59

Figure 5.1 Schematic representation of the depositional history 65

Figure 5.2.2-1 Schematic cross sections of the lower portion of the Chinle Fm.. 74

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CHAPTER I

INTRODUCTION

1.1 Purpose

The purposes of this study are to 1) place the lower part of the Chinle

Formation of central Utah into a regional depositional framework; 2) place the strata

into sequence stratigraphic context; and 3) constrain the tectonic and climatic effects

on the Late Triassic landscape. The recognition and correlation of three regional

sequence boundaries provide a temporal framework for the various depositional

facies associated with each aggradational sequence. Paleoclimatic and paleoecologic

indicators recorded within Chinle paleosols and lacustrine deposits provide

independent data to support conclusions drawn regarding possible forcing

mechanisms driving Chinle deposition. The specific objectives of this study are as

follows:

1) Describe the sedimentology, depositional facies, and large scale architecture

of the Chinle Formation exposed in south central Utah.

2) Use sequence stratigraphic methods to create a temporal framework for these

strata.

3) Collect paleoecologic data and analyze paleoclimatic indicators both within

the deposits and the paleosols in order to provide an independent, temporally

continuous paleoecologic and paleoclimatic record.

4) Interpret a detailed depositional history for these strata that includes probable

forcing mechanisms controlling Late Triassic landscape evolution.

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5) Clarify traditional lithostratigraphic nomenclature and present evidence

supporting possible temporal correlations among various lithomembers.

1.2 Location

The Chinle Formation is present in outcrop and in the subsurface throughout

much of the southwestern United States (Figure 1.2-1). This study focuses on

extensive exposures of Triassic strata exhumed in large uplifts (San Rafael Swell,

Waterpocket Fold, and Monument Valley Upwarp) associated with the Laramide

orogeny (Stewart et al., 1972). Four distinct field areas include the San Rafael Swell

south of I-70, the northern part of Capitol Reef National Park, the Circle Cliffs area

within Grand Staircase Escalante National Park, and the White Canyon, Red Canyon,

and North Wash areas within Glen Canyon National Recreation Area. Geophysical

well logs were obtained from adjacent areas where the Chinle is present in the

subsurface. The total field area consists of over 5,000 square miles of remote central

Utah desert (Figure 1.2-2). Access to many of the measured sections requires four-

wheel drive vehicles, mountain bikes, or long hikes over steep terrain. Scientific

permits were obtained from Grand Staircase Escalante NP, Capitol Reef NP, and

Glen Canyon NRA.

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Figure 1.2-1 Distribution of Triassic outcrops in the southwestern United States (from

Dubiel, 1994).

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Figure 1.2-2 Field area map of central Utah depicting location of measured sections and

well logs. SWC = Spotted Wolf Canyon, TFC = Three Fingers Canyon, EC =

Eardley Canyon, TM = Temple Mountain, CC = Chute Canyon, CL = Chutes

and Ladders, BS = Blueberry Spring, BC = Bell Canyon, ER = End of the Road,

HS = Hidden Splendor, PT = Pasture Track, NW = North Wash, CP = Copper

Point, BN = Blue Notch, JC = Jacobs Chair, FC = Fry Canyon, MP = Muley

Twist, CC = Long Canyon Overlook, FL = Fish Lake, PB = Park Boundary, CR

= Chimney Rock.

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1.3 Methods

1.3.1 Field Methods

During the summer of 2004 and the winter and spring of 2005, 27 sections

were measured, described, and sampled. Sections were measured with a jacob staff

and Brunton compass. With the exception of several sections measured in the San

Rafael Swell, Triassic strata were found to dip less than 10 degrees from horizontal.

Thicknesses of several highly-resistant cliffs were determined using digital imaging

techniques. Measurements start at the Moenkopi-Chinle contact and stop at a

regionally recognizable paleosol at the base of the Petrified Forest Member. Section

locations were selected on the basis of accessibility (steepness) and exposure quality.

Where permissible, moderate hand trenching was used to increase the amount and

quality of exposure, especially in frothy, weathered slopes. Hand samples of

sandstones and paleosols were collected from many of the sections for petrographic

analyses. Numerous digital photographs were taken at each of the measured section

localities. In addition to these photographs, many photopans were taken from

appropriate vantage points across the field area in order to help constrain the meso-

scale architecture of the paleovalley fill. Paleocurrent data were collected in the field

and analyzed using the computer program GEOrient 9.0 to aid paleogeographic

reconstruction.

1.3.2 Petrographic Methods

Hand samples of sandstones from the Shinarump, Monitor Butte, Temple

Mountain, and Moss Back Members were collected from 13 measured section

locations listed in Appendix III. Hand samples were made into thin sections and

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stained for potassium feldspar. Three hundred points were counted in a 1 mm grid for

each of the 35 thin sections. Ten grain types were identified for each sample and the

percentage of the following three derived variables: total quartz, total feldspar, lithic

and unstable constituents was computed (Figure AIII.1).

1.3.3 Subsurface Methods

Digitized geophysical well log data were obtained for 21 wells selected from

areas adjacent to field locations where the Chinle is found in the subsurface. Surfaces

were correlated between 18 of the logs using sequence stratigraphic methods

described by Posamentier and Allen (1999) and Van Wagoner et al. (1990). The top

of the Moss Back Member was selected as a datum because of its flat upper surface

and lateral continuity across the field area (Dubiel, 1983). This surface, marked in

outcrop by either a transition from a resistant, well-sorted sandstone to deposits with a

much lower sandstone/mudstone ratio or by an intensely pedogenically modified

horizon, is readily observable in the subsurface.

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CHAPTER II

GEOLOGIC SETTING

The Late Paleozoic to mid-Mesozoic geologic record preserved in central

Utah is both spatially and temporally complicated by several orogenic events,

laterally diverse paleolandscapes, changes in relative sea level, and shifting climatic

regimes. Despite the interplay between 1) tectonics 2) stratigraphy 3) paleogeography

4) paleoclimate and 5) paleoecology, the wealth of research conducted during the past

century by numerous workers warrants the organization of the information into the

aforementioned subheadings in order to present a tangible synopsis of the geologic

events and environmental conditions surrounding the deposition of the Chinle

Formation.

2.1 Tectonics

The geologic history of Utah is broken down into eight major phases, the first

three of which precede the Triassic (Hintze, 1988). Phase I includes the Archean

craton found in the northern third of the state and the Early-Proterozioc metamorphic

terrane which was added to the craton in the Mid-Proterozic and now comprises the

basement rock for the remainder of the state. Phase II was dominated by state-wide

deposition along a west facing passive margin (Hintze, 1988). Collisional tectonics

during Phase III associated with the Antler Orogeny in the northwest and the

Ouachita Orogeny to the southwest shifted the locus of deposition during this time to

the respective Oquirrh and Paradox Basins. Collisional tectonics and subduction

along the west coast of Pangaea occurred from the Triassic through the Eocene during

Phases IV (Sonoma and Nevada Orogenies), V (Sevier Orogeny), and VI (Laramide

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Orogeny). Large scale Oligocene and Miocene volcanism and later uplift and basin-

range extension occurred during the most recent Phase VII and VIII respectively.

Of these ever-shifting tectonic regimes, the assemblage of Pangaea during the

Late Paleozoic and Early Mesozoic (Phases III and IV) most strongly affected the

character of the Late Triassic landscape and the nature of sediments from that time

period. Prior to the assemblage of Pangaea, central Utah was located on the western

passive margin of Laurussia. Beginning in the Carboniferous, Gondwana collided

with Laurussia from the south resulting in the initial stage of the assemblage of the

supercontinent Pangaea. Regionally associated with the collisional event, the

Carboniferous to Permian Ouachita orogeny resulted in the uplift of the Ancestral

Rockies stretching from the modern areas of Price, Utah, to Grand Junction,

Colorado, to Santa Fe, New Mexico (Prothero and Dott, 2004). Deposition occurred

in the Paradox Basin adjacent to these uplifts (Condon, 1997).

The initiation of the Antler and Sonoma orogenies during the Devonian and

Late Permian, respectively, is marked by a transition along the western margin of

Pangaea from a passive margin to an actively subducting convergent boundary

(Figure 2.1-1). Sonoman subduction continued through the Late Triassic when the

western flank of Pangaea was occupied by an island arc stretching northward from

present day California and a continental arc that wrapped around to the south into

present day Mexico.

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Figure 2.1-1 Late Triassic tectonic setting of the Colorado Plateau after Dubiel, 1994,

Lawton, 1994, and Riggs et al., 1996.

Many workers (Dickinson, 1981, Blakey and Gubitosa, 1983, Dubiel et al.,

1991, Dubiel, 1994) classify the Chinle depositional basin as a back-arc basin

associated with Sonoma-related subduction along the west coast of Pangaea. Lawton

(1994) classifies the Chinle depositional basin as a back-bulge basin associated with

the subduction of the pre-Farallon slab. In both models, subsidence of the basin is

attributed to dynamic topography resulting from viscous flow in the mantle above the

subducting slab. However, Lawtons model places a dynamic forebulge in central

Nevada roughly 400 km east of the subduction zone during both Early and Late

Triassic time. The Chinle depositional basin was flanked to the east by the Ancestral

Rockies, which remained a topographic high until it was overlapped in the Jurassic by

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the Morrison Formation, and to the south by the continental portion of the arc system

often referred to as the Mogollon Highlands.

The breakup of Pangaea started in the Late Triassic with rifting of the proto-

Atlantic margin. Thermal uplift associated with the rifting affected deposition of

correlative Triassic sediments in the Dockum Basin of Texas. However, the effect of

these tectonic events on the Chinle Basin is unclear (McGowen et al., 1979, Dubiel et

al., 1991, Woody, 2003).

Salt tectonics have been shown to have affected deposition of the Chinle east

of the field area in the Paradox Basin (Blakey and Gubitosa, 1983, Hazel, 1991,

Dubiel, 1994, Condon, 1997). Halokinesis was limited laterally by the extent of the

Pennsylvanian Paradox Formation and likely had little direct effect upon deposition

of the Chinle further west in the field area.

Deep-seated deformation, associated with the shallowing of the subducting

Farralon Plate during Late Cretaceous-Eocene Laramide orogeny, resulted in the

formation of several large fault propagation folds, which include the Waterpocket

Fold, Monument Upwarp, San Rafael Swell, and the Uinta Dome. Modern exposures

of Triassic strata are exhumed in these uplifts. Erosion resulting from Late Cenozoic

thermally related uplift, combined with modern aridity, provides excellent exposure

of previously buried Triassic strata across the Colorado Plateau.

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2.2 Stratigraphy

Figure 2.2.2-1 Nomenclature and gross stratigraphic architecture of Late Paleozoic to Mid-

Mesozoic strata in central Utah.

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2.2.1 Pennsylvanian -Permian

Deposition during the Appalachian-Ouachita orogeny occurred in the Paradox

foreland basin adjacent to the Ancestral Rockies (Hintze, 1988) where deposition of

Pennsylvanian marine strata preceded the deposition of primarily continental Permian

strata. The Permian Cutler Group was deposited as siliciclastic sediment prograded

southwestward off the Uncompahgre Highland (the Ancestral Rockies) (Doelling et

al., 2000). Proximal to the Uncompahgre, in the deepest part of the basin, the Cutler

Group is termed the Cutler Formation, undivided (Condon, 1997). The Cutler

Formation, undivided grades distally into lithologically distinct units which include in

ascending order the lower Cutler beds, Cedar Mesa Sandstone, Organ Rock Shale,

DeChelly Sandstone, and the White Rim Sandstone (Figure 2.2.2-1) (Condon, 1997).

Shallow marine carbonate was deposited during this time period roughly 160

kilometers further southwest from the hinterland in the Grand Canyon region

(Coconino Limestone, Toroweap Formation). A major marine transgression during

the Late Permian extended carbonate deposition northward, represented by the

Kaibab and Black Box Limestones in the northern portion of the field area.

2.2.2 Early Triassic

The Permo-Triassic boundary is marked by the Tr-1 unconformity, which

separates Permian strata from the overlying Early Triassic Moenkopi Formation

(Figure 2.2.2-1) (Pipringos and OSullivan, 1978). In southwestern Utah the

unconformity at the base of the Moenkopi is characterized by as much as 180 meters

of karstification of the Kaibab Limestone (Jensen, 1984). Erosional truncation of the

Kaibab marks the unconformity in the San Rafael Swell, Capitol Reef, and Circle

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Cliffs area, however Stewart et al. (1972) suggest the contact becomes conformable

east of the field area.

The Hoskinnini and Black Dragon Members mark the base of the Moenkopi

Formation. These two members consist of terrestrially deposited, wavy-bedded

sandstones, siltstones, evaporites, and coarse chert pebble conglomerates that fine

upward to marine sandstones, shales, and minor carbonates (Huntoon et al., 2000,

2002, Mitchell, 1985). The thickness of these two members is inversely proportional

to the thickness of the underlying White Rim Sandstone because these deposits onlap

the paleorelief formed during the Tr-1 unconformity.

The Black Dragon grades upward and intertongues with the overlying Sinbad

Limestone (Mitchell, 1985). Lithologic textures, which include features such as

cross-bedded ooids, combined with subsurface data suggest deposition in a shallow

marine system with a low depositional slope (Mitchell, 1985). The Sinbad is present

in the San Rafael Swell and is thickest at Capitol Reef National Park but pinches out

to the south and east and is neither present at Muley Twist nor at Hite (Stewart et al.,

1972, Mitchell, 1985). The Black Dragon and Sinbad were deposited during the

Smithian transgression which encroached from the northwest (Dubiel, 1994).

The Sinbad grades and intertongues with the ledgy fine-grained sandstone,

siltstone, and mudstones that comprise the Torrey and Moody Canyon Members. The

Torrey thickens proximally to the east and represents the shoreward equivalent of the

Virgin Limestone Member exposed west of the field area in south-central Utah

(Mitchell, 1985). Both the Torrey and overlying Moody Canyon Members exhibit

very laterally continuous beds, which are identifiable both in the field and in

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subsurface data. Blakey (1974) identified four facies within the Torrey Member,

which he attributes to deltaic deposition, while Mitchell (1985) and Reif and Slatt

(1979) believe deposition occurred within a coastal tidal flat environment. The

Moody Canyon Member is slightly finer-grained overall than the Torrey Member and

represents deposition either in a restricted marine (Mitchell, 1985) or clastic-

dominated sabkha (Blakey, 1974) environment. The Moody Canyon marks the end of

Early Triassic deposition during the late Spathian Regression (Dubiel, 1994).

2.2.3 Late Triassic

The Tr-3 unconformity separates the Early Triassic Moenkopi Formation from

the Late Triassic Chinle Formation (Figure 2.2.2-1). Fluvial incision and interfluve

pedogenesis of the underlying Moody Canyon Member marks the base of the Chinle

Formation, which is characterized by mudstones, sandstones, limestones, and

conglomerates. The depositional facies represented in the Chinle include a variety of

fluvial, fluvio-deltaic, lacustrine, paludal, and minor eolian deposits, all representing

deposition in an entirely continental setting.

Early paleontologic work by Gregory (1917) placed the Chinle Formation in

the Late Triassic. Ar/Ar dates on multi-grain zircon fractions provided a maximum

age for the Black Forest bed in Petrified Forest National Park (PEFO) of 207 2 Ma.

(Riggs et al., 1994). U-Pb dates of zircons from the same stratigraphic interval

indicate a maximum age of 209 5 Ma. (Riggs, 2003). Sequence stratigraphic

correlations do not extend as far south as PEFO in this study, however preliminary

field data (Woody, 2003, Demko, pers commun.) suggests both the Sonsela and the

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Black Forest Bed are time-correlative to some portion of the Petrified Forest Member

of central Utah.

Ash (1980, 1987) identifies three Late Triassic floral zones in the Chinle

Formation: 1) middle Carnian Eoginkgoites zone; 2) late Carnian Dinophyton zone;

3) Norian Sanmiguelia zone (Figure 2.2.3-1). Ash finds the Eoginkgoites flora in the

Shinarump and Temple Mountain members, Dinophyton in the Monitor Butte and the

lower portion of the Petrified Forest Member, and Sanmiguelia in the upper portion of

the Petrified Forest Member, and the Owl Rock and Church Rock Members (Ash

1975, 1976, 1977, 1980, 1987). Litwin et al. (1991) divide the Chinle into three

palynofloral zones which correspond almost identically to the megafossil plant zones

of (Ash, 1980, 1987).

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Figure 2.2.3-1 Triassic timescale modified from IUGS (2004).

The lithostratigraphy of the Chinle Formation contains a myriad of localized

members identified on the basis of lithologic characteristics and stratigraphic

relationships that have proven difficult to correlate across the entire Chinle basin.

Lucas (1993) proposed basin-wide correlations of Chinle and adjacent Late Triassic

strata that ignored the complex spatial variability of the Late Triassic landscape, lack

accurate intraformational temporal control, and confuse the preexisting, well-defined

member nomenclature (Dubiel, 1994). The author agrees with the criticisms of Dubiel

(1994), and Woody (2003), and will use the existing nomenclature summarized in

Stewart et al. (1972) when referring to large scale lithostratigraphic trends. However,

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intraformational correlations will be based upon chronostratigraphically-significant

surfaces, which, by definition, may cross these lithologic boundaries. The sub-

formation units recognized in the study region include the Shinarump, Temple

Mountain, Monitor Butte, Moss Back, Petrified Forest, Owl Rock, and Church Rock

Members.

Shinarump Member. The Shinarump Member of the Chinle Formation is

recognized as a resistant, quartz rich, light colored sandstone and conglomerate

usually less than 10 meters thick. The Shinarump has been interpreted to have been

deposited by braided streams in broad thin sheets of interconnected sandstone

bodies (Blakey and Gubitosa, 1984). The Shinarump Member rests unconformably

upon the Moenkopi Formation in incised valleys as large as 16 kilometers wide and

55 meters deep (Stewart et al., 1972). The Shinarump represents the first of a series of

deposits that fill these large paleovalleys. In all cases within the study area, the

Shinarump occupies the basal portion of the Painted Desert paleovalley (Figure 2.2.3-

2) and is in direct contact with the underlying Moenkopi Formation (Blakey, 1983,

Dubiel et al., 1999). In some cases the Shinarump is found on top of the Moenkopi

proper, while in other places it is found overlying a pedogenically modified

Moenkopi Formation described as mottled strata (Stewart et al., 1972). The

mottled strata, as described by Stewart et al. (1972), includes several pedogenic

horizons, not only those formed in Moenkopi parent material, but also younger, well

developed paleosols that formed in stratigraphically higher Chinle deposits.

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Figure 2.2.3-2 Map depicting field area (as depicted in Fig 1.2-2) in relationship to the Painted

Desert Paleovalley (modified from Dubiel et al., 1999).

Stewart et al. (1972) describe the contact between the Shinarump and

overlying strata as conformable stating, the Shinarump Member in most areas grades

upward into and intertongues with claystone, siltstone, or clayey sandstone.

However, Stewart et al. (1972) also include a dramatic photograph (Figure 7 in

Stewart et al., 1972) which illustrates nearly 13 meters of truncation visible in a single

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outcrop. Demko (2003) interprets the upper contact as a regionally extensive

sequence-bounding unconformity.

Temple Mountain Member. The Temple Mountain Member of the Chinle

Formation was first described in the San Rafael Swell by Robeck (1956) as a purple

and white mottled unit, dominated by siltstones and mudstones with minor

conglomeratic sandstone lenses, found between the Moenkopi Formation and

overlying Monitor Butte and Moss Back Members. Near Hidden Splendor Mine

(Figure 1.2-2), southwest of the type section at Temple Mountain, the unit has been

interpreted to either underlie (Robeck, 1956, Stewart et al., 1972) or intergrade

(Pavlak, 1979) with the Monitor Butte Member. The Temple Mountain reaches a

maximaum thickness of 33 meters north of the type section; further north, the entire

unit is truncated by the overlying Moss Back Member (Robeck, 1956). Stewart et al.

(1972) describe the Temple Mountain as part of the ubiquitous mottled strata based

upon its characteristic color mottling and its basal stratigraphic position within the

Chinle Formation.

Monitor Butte Member. The Monitor Butte Member is made up of green-

gray mudstones and siltstones, thin sandstones, and occasional thin coals and

limestones representing deposition in fluvial, lacustrine, lacustrine-deltaic, and

paludal depositional settings (Dubiel, 1985, Demko, 2003). The majority of the

Monitor Butte in the field area consists of an overall coarsening-upward succession

comprised of sandy limestones and organic rich mudstones, thinly bedded micaceous

sandstones, and burrowed, pedogenically modified sandstones and mudstones

interpreted to represent prodelta, delta front, and distributary facies within a

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progradational fluvio-deltaic lacustrine system (Dubiel, 1985). Deformational

features including contorted, slumped, and steeply dipping bedding have been

attributed to syndepositional mud diapirism (Stewart et al., 1972, Dubiel, 1985).

The Monitor Butte Member is recognized throughout much of southeastern

Utah in White Canyon, Capitol Reef, Circle Cliffs, and in the southwestern portion of

the San Rafael Swell (Figure 1.2-2). The Monitor Butte overlies either the Moenkopi

Formation or the Shinarump Member where present. Stewart et. al. (1972) state the

Monitor Butte either interfingers with the Shinarump Member or is separated by an

unconformity depending upon location within the basin. Dubiel (1983) claims the

Shinarump interfingers with the lowermost organic rich mudstones of the Monitor

Butte Member in White Canyon. Demko (2003) has interpreted the boundary at the

base of the Monitor Butte to represent a flooding surface sequence boundary (FSSB),

followed by deposition of the Monitor Butte in a highstand systems tract. In Demkos

(2003) model, the boundary between the Monitor Butte and the overlying Moss Back

Member also represents a sequence boundary.

Moss Back Member. The Moss Back Member is a brown, fine- to coarse-

grained cliff-forming conglomeratic sandstone characterized by planar tabular,

overturned, and trough crossbedding and planar bedding (Dubiel, 1983, Stewart et al.,

1972). The Moss Back is exposed as either multistorey sheet or channelform sand

bodies. Dubiel (1983) and Blakey and Gubitosa (1984) interpret deposition of the

Moss Back Member by braided streams based on high sandstone/mudstone ratios,

extensive planar-tabular cross beds, and the multistorey broad sheet-like architecture.

Pavlack (1979) identifies well defined fining-upward sequences and lateral accretion

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surfaces in the Moss Back in the San Rafael Swell, and interprets deposition by

braided as well as high- to low-sinuosity streams.

The Moss Back averages 20 meters thick, but individual channels can vary

from over 35 meters to zero where they pinch out at the channel margins (Stewart et

al., 1972, Demko, 2003). The base of the Moss Back is marked by an erosional

surface that truncates the Monitor Butte Member, Temple Mountain Member, and

Moenkopi Formation. The base of the Moss Back has been interpreted as a sequence

boundary marking an unconformable basinward-shift in facies (Demko, 2003). The

top of the Moss Back is conformable with overlying Petrified Forest Member strata.

The transition, while conformable, is marked by a lithologic change from a resistant

brown sandstone marking the top of the Moss Back to a regionally identifiable red

aggradational paleosol located at the base of the Petrified Forest Member.

Petrified Forest, Owl Rock, and Church Rock Members. These three

members are dominated by fluvial, lacustrine, and eolian deposits respectively located

stratigraphically above the aforementioned members. The Petrified Forest Member is

the first of the Chinle deposits to overlap the Painted Desert paleovalley initially cut

by the Tr-3 unconformity. The transition between the lower deposits making up the

incised valley fill and the overlying, relatively unrestricted deposits is marked by a

change in fluvial style from large confined stream deposits of the Moss Back to

laterally migrating streams, and by slower vertical sediment accumulation marked by

increased pedogenesis. Demko (2003) and Heckert (2004) recognize a regional

unconformity at the top of the Petrified Forest Member that reflects a non-

depositional hiatus between the Petrified Forest Member and the conformable Owl

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Rock and Church Rock Members. While recognition of the upper members of the

Chinle is vital to the success of this study, detailed analysis of these members is

beyond the scope of this project.

2.2.4 Jurassic

The Chinle is separated from the overlying Early Jurassic Glen Canyon Group

(Wingate Sandstone, Kayenta Formation, and Navajo Sandstone) by the J-0

unconformity (Pipringos and OSullivan, 1978). The unconformity leaves 40 meters

of the Church Rock Member of the Chinle Formation in Red Canyon (Stewart et al.,

1972) but truncates the entire member at Ward Terrace in north central Arizona

where Wingate-correlative Moenave Formation unconformably overlies the Owl

Rock Member. The Wingate Sandstone forms distinct red vertical exposures over 65

meters thick comprised of large-scale trough crossbedded eolian sandstone. Overlying

the Wingate, fluvial-eolian deposits of the Kayenta Formation crop out as

recognizable banded cliffs and ledges. The Navajo Sandstone unconformably overlies

the Kayenta and is comprised of light colored well sorted large-scale trough-

crossbedded sandstone deposited in a giant Jurassic aged erg that stretched from

present day Arizona to Montana. These three formations serve as a roughly 500

meter thick resistant caprock overlying Triassic rocks throughout the field area and

make up the bulk of the scarps in the San Rafael Reef, Capitol Reef, and Glen

Canyon field areas.

2.3 Paleogeography

Pangaea was initially formed by the collision of Laurussia and Gondwana in

the Carboniferous (see section 2.1). By the Triassic, much of Asia had been added to

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the eastern portion of the supercontinent which now spanned from 85 degrees N to 90

degrees S latitude (Parrish, 1993). During this time, central Utah was located at 10

degrees N latitude on the western margin of the supercontinent (Figure 2.3-1) (Van

der Voo et al., 1976, Pindell and Dewey, 1982). Low eustatic sea level during the

Late Triassic (Haq et al., 1987, Vail et al., 1977) placed central Utah approximately

500 km from the Panthalassic coast (Dubiel, 1994, Riggs et al., 1996).

Figure 2.3-1 Reconstruction of Triassic Pangaea showing location of field area. Modified

from http://web.uvic.ca/~rdewey/eos110/webimages.html, 2005.

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Studies of paleocurrent and provenance (Riggs, 1996, Stewart et al., 1986,

Stewart, 1972, Albee, 1957, Gehrels, 1993) suggest the Chinle fluvial systems flowed

west and northwest from as far west as the Dockum Basin in Texas. The

Uncompahgre and the Ancestral Front Range acted as the major sediment source for

the Chinle deposited in the Eagle, Piance, and Uinta Basins in northeastern Utah and

northwestern Colorado (Dubiel, 1991). The source areas for the Chinle of central

Utah were the Uncompahgre to the east and the active continental volcanic arc to the

south near the Mexican border (Dubiel, 1991, Pavlack, 1979). Chinle deposits further

south in PEFO were mainly sourced from the southern and southwestern continental

arc (Schultz, 1963, Stewart et al., 1972, Stewart et al., 1986, Riggs et al., 1994,

Woody, 2003).

Paleogeographic reconstructions depicting the evolution of the Late Triassic

landscape which reflect paleocurrent and provenance data, the distribution of

interpreted depositional settings, and evidence of increasing aridity through Chinle

time (see section 2.4) are shown in Figure 2.3-2. The Shinarump through Petrified

Forest Members are shown to represent fluvial and fluvio-lacustrine drainage

networks draining northwestward toward the Late Triassic shoreline located in central

Nevada. Deposition of the younger Owl Rock and Church Rock Members is

interpreted to have occurred within a more arid, internally drained basin dominated

by ephemeral fluvial, palustrine, lacustrine, and eolian deposition.

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Figure 2.3-2 Paleogeographic reconstructions of Late Triassic Chinle depositional systems

(from Blakey, http://jan.ucc.nau.edu/~rcb7/RCB.html, 2003).

2.4 Paleoclimate

The Triassic climate is interesting because the large Pangaean landmass had a

dramatic effect upon global atmospheric and oceanic circulation. Worldwide, the

Triassic is marked by areally extensive evaporite and red bed deposition suggestive of

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arid conditions. However, poorly drained paleosols, perennial fluvial and lacustrine

deposits, and a diverse assemblage of preserved flora from the Chinle Formation are

suggestive of more humid conditions for western subequatorial Pangaea. Not only

does the Chinle basin appear to be wetter than the majority of other Triassic

depositional environments, it is characterized by sedimentary paleoclimatic indicators

that suggest more humid conditions than those of the underlying Permian Cutler

Group and the overlying Jurassic Glen Canyon Group eolianites. Using the then

newly developed ideas of plate tectonics, Irving and Briden (1964) used paleolatitude

data to reconstruct Pangaean paleoclimate and found that the Triassic lacked

latitudinal climatic zones typical of todays global climate. In 1973, Robinson further

described the climate of Pangaea as monsoonal. Later workers have described the

mechanisms driving monsoonal circulation (Kutzbach, 1987) and have described

paleoclimatic indicators preserved in the Chinle which are suggestive of overall warm

temperatures and episodically fluctuating moisture conditions (Dubiel et al., 1991,

Parrish, 1993).

Triassic monsoons are attributed to the northward movement of Pangaea

beginning in the Late Paleozoic. During this time the large continental landmass

became centered around the equator (Dubiel et al., 1991) and differential heating of

the northern hemisphere during the summer created a large low pressure system and

cross-equatorial circulation patterns (southerly winds) dominated. During the

northern hemisphere winter, the low pressure shifted southward resulting in northerly

wind patterns. The center of the continent experienced arid conditions, however,

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monsoonal circulation would have been strong enough during the Triassic to draw

moisture from Panthalassic Ocean eastward to the Chinle basin.

An abundance of data outlined by Dubiel et al. (1991) and Parrish (1993)

suggests the Chinle basin experienced alternating wet/dry conditions typical of a

monsoonal climatic setting. Numerous vertisols characterized by well-developed

slickensides, desiccation cracks, and occasionally apparent gilgai are found in the

Temple Mountain, Monitor Butte, and Petrified Forest Members of the Chinle.

Optimal conditions for vertisol development are subtropical to tropical monsoonal

climates that undergo alternate dry and wet periods (Blodgett, 1985). The majority of

modern vertisols are found in areas affected by the Indian and East African monsoon

(Brady and Weil, 2004).

Another indication of fluctuating soil hydrology are color mottles. Mottles

form when iron and magnesium in the soils change valence and mineral form to

adjust to various redox conditions in the soil (Montgomery et al. 2000). Paleosols

exhibiting mottles of both low chroma colors like green and gray (typical of reducing

conditions) and high chroma colors like purple and red (typical of oxidizing

conditions) are typical of Temple Mountain and Monitor Butte paleosols found near

the base of the Chinle. Evidence of seasonality is also seen in growth bands of

unionid bivalves where seasonally wet conditions lead to turbid, high suspended load

flow conditions in which the organism stops secreting carbonate and a dark band is

preserved in the shell (Dubiel et al., 1991).

Indicators of seasonality persist through Chinle time; however, minor eolian

deposits and increased evidence of desiccation in the Church Rock Member suggests

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overall climatic drying through the Late Triassic (Blakey and Gubitosa, 1983, Dubiel

et al., 1991). Pangaea moved further north and had begun to break apart by the

Jurassic. The climate of the region became fully arid by this time as indicated by the

eolian erg deposits of the Wingate Sandstone.

Although recent workers agree the climate of the Chinle depositional basin

was dominated by monsoonal wet/dry seasonality that became more arid through

time, a disparity exists within the literature concerning how much rainfall the basin

received during the deposition of the lower portion of the Chinle. Ash (1972, 86),

Creber and Ash (1990), and Ash and Creber (1992) interpret a humid climate based

on their interpretation of Chinle floral assemblages. Aquatic vertebrate assemblages,

which include phytosaurs and amphibians, also indicate at least local areas with

perenial surface water. Demko et al. (1998) suggest the plant and vertebrate fossil

assemblages preserved in the lower portion of the Chinle may be taphonomically

biased. According to Demko et al. (1998) abundant plant fossils are constrained to

areas located within Chinle paleovalleys where locally high water tables are suitable

for the preservation of organic material resulting in a preservational bias of more

humid paleontologic indicators.

Pollen analysis by Gottsfield (1972) and faunal (Parrish et al., 1986) and floral

(Ash, 1999) specimens representative of upland environments have been interpreted

as indicative of more arid conditions. Evidence of extended dry periods can be seen in

crayfish burrows which exhibit multiple aestivation chambers created in order to

follow falling water tables (Hasiotis and Mitchell, 1989, Hasiotis et al., 1989,

Kowaleski and Demko, 1996). Paleosol analyses by (Dubiel, 1994, Hasiotis and

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Mitchell, 1989) support the interpretation of dry to hydrologically well-drained

conditions.

Despite temporal and spatial variability and the natural biases associated with

a multitude of climatic proxies, the climatic setting of the Chinle depositional basin

can be described as generally warm, evolving from semi-arid to arid conditions,

punctuated by seasonal variations in precipitation resulting from cross-equatorial

circulation patterns associated with the Pangaean monsoon.

2.5 Paleoecology

Not only was the Chinle deposited during an interesting paleoclimatic

interval, it also coincides with a critical period of biologic evolution. The Chinle

Formation is paleontologically significant because it preserves a record of the

continuing diversification of terrestrial life following the largest extinction event in

Earths history, the Permo-Triassic extinction. The Triassic marks the first appearance

of mammals and the diversification and ascension of dinosaurs. Numerous studies

(Ash, 1972, Gottesfeld, 1972, Ash, 1975, Ash, 1986, Ash, 1987, Ash and Creber,

1992, Ash, 1999) of outstandingly preserved vertebrate and plant fossils preserved

within the Chinle Formation (especially in Petrified Forest National Park) provide

detailed insights into the paleoecology of subequatorial Pangaea. In addition,

continental ichnology has become an increasingly important tool used to identify

hidden biodiversity preserved in paleosols of the Chinle Formation. Trace fossils have

been successfully used as paleoecologic indicators of soil chemistry and hydrology

(Hasiotis and Mitchell, 1989, Hasiotis, 2002, Hasiotis et al., 2002, Hasiotis, 2004).

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CHAPTER III

DATA AND RESULTS

3.1 Unit Descriptions - Lithology, Facies, and Architecture

Descriptions of the lower portion of the Chinle Formation are organized in

this chapter on the basis of the established lithostratigraphic nomenclature outlined in

section 2.2.3. The variable occurrence of the ubiquitous mottled strata and the

poorly defined relationship between the Monitor Butte and Temple Mountain

Members, in addition to disparities between sequence stratigraphic and

lithostratigraphic correlations, makes such organization difficult. As a result, several

smaller units are described individually outside of the lithostratigraphic classification

scheme in order to minimize any confusion and to aid the transition to the next

chapter.

3.1.1 Basal Unconformity

The Tr-3 unconformity at the base of the Chinle Formation is marked by

erosional truncation of the Moenkopi Formation (measured sections 2, 3, 5, 6, 9, 11,

12, 14-17, 20, 23-27) and/or extensive pedogenesis of those strata (measured sections

1, 4, 6, 7, 10, 13, 18, 19, 21, 22). The contact varies from extremely sharp, where

Late Triassic Strata directly onlap truncated Early Triassic rocks, to gradational where

the uppermost Moenkopi exhibits gleyed joints along bedding surfaces and also

where the occasional penetration of crayfish burrows renders the contact variable

(Figure AI.56).

The paleosols formed in the Moenkopi parent material range from 1-7 meters

thick and are characterized by large color mottles often dark purple, gold, and black

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haloed by light grey in an overall light purple background (Figure AI.8). Grain-size of

these paleosols is variable, ranging from medium grained sand to dominantly silt and

clay. In most cases, especially where the paleosol is finer grained, these rocks weather

to form resistant, light-colored ledges (Figures 3.1.1-1, AI.12, AI.41). These ledges

extend laterally for several kilometers exhibiting gentle paleorelief of up to 10 meters.

Evidence of soil fabrics include large root traces and burrows; however, often the

identification of such features is difficult because the rocks are extensively

bioturbated and have undergone multiple stages of oxidation/reduction (Figure AI.8).

Evidence of horizonation in these paleosols is limited to gradational changes in color,

color-mottle size and orientation, and grain size Figure (AI.56).

Figure 3.1.1-1 Laterally extensive unconformity paleosol (B) found at the base of the Chinle

Formation east of MS 27. Notice overlying left-to-right dipping lacustrine

clinoforms (A) in the Monitor Butte Member.

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These paleosols are classified as gleyed oxisols (after Mack et al., 1993).

While the most striking feature of these soils is the various color mottles, the most

prominent process is extensive weathering. The closest modern analogs for these soils

are ultisols formed in tropical regions marked by variable water tables. These soils,

known as plinthites or laterites, are thick, iron-rich, humus poor, and contain color

mottles and clays which become very hard when dried (Brady and Weil, 2004). The

well-indurated nature of similar, Quaternary-aged, paleosols identified in monsoonal

regions of present day Australia, has been shown to affect fluvial geomorphology by

limiting both vertical and lateral channel adjustment (Nanson et al., 2005).

3.1.2 Shinarump Member

The first Late-Triassic rocks to be deposited after the Tr-3 unconformity

include the coarse-grained sand and gravel deposits of the Shinarump Member. These

strata are recognized at the base of the Chinle Formation where the lower portion of

the Chinle is the thickest (measured sections 17, 19-21, 25-27). In the field area the

Shinarump varies from 30 (Figure AI.76) to zero meters thick where it both onlaps

the paleovalley margin and is truncated by overlying units (Figure AI.82). The basal

contact of the unit is marked by sharp truncation of the Moenkopi Formation at all

localities within the field area. Often the upper meter of Moenkopi strata below the

contact is gleyed. The gleyed strata exhibit light grey color bands several centimeters

thick, parallel horizontal bedding, and occasional vertical joints and contrast sharply

with the overall red color of the Moenkopi Formation.

The Shinarump generally consists of yellow-white, coarse, quartz sand and

gravel conglomerate. Sand and pebble-sized sediments often exhibit repeating,

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decimeter-scale trough-cross beds while the coarser fraction consists of smaller-scale

planar or massive bedding. In addition to these coarse-grained sediments, thinly

laminated lenses less than 3 meters thick, consisting of fine-grained sand and silt,

were found in Red Canyon and near Chimney Rock. In both cases these fine-grained

deposits contained abundant plant fossils and minor amounts of root traces and

desiccation cracks.

On a larger scale, the Shinarump is comprised of several storeys of

amalgamated tablular sand sheets interspersed with smaller lenticular-shaped

channels. Since these channelforms fill a wide, low relief, paleovalley, the resultant

sandbody geometry for the Shinarump is a broad lens.

3.1.3 Post-Shinarump Unconformity

The upper contact of the Shinarump is marked by either erosional truncation

or by a deeply weathered paleosol. Most often the gentle slope-forming nature of the

overlying strata (Monitor Butte Member) obscures the nature of this contact;

however, several meters of truncation of the Shinarump can be seen in the excellent

exposures near the Fish Lake and Chimney Rock measured sections in Capitol Reef

National Park (Figure AI.82).

A well-developed gleyed oxisol, similar in nature to the paleosol described in

section 3.1.1, marks the base of the Monitor Butte Member immediately above the

coarse-grained deposits of the Shinarump Member in White Canyon, at Jacobs Chair

(MS 19), and in Red Canyon (MS 21). At Jacobs Chair, the paleosol is fine-grained

and nearly 6 meters thick. Color mottles found in this paleosol range from dark

purple, red, grey, gold, white, green, to dark black in color (Figure AI.59). Remnant

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sedimentologic horizonation is crude, however pedogenic indicators such as crayfish

burrows and rooted horizons indicate slow overall soil accumulation under highly

fluctuating hydrologic conditions. At Red Canyon the paleosol is much thinner. Here

the upper 5-30 cm of the coarse-grained Shinarump has been altered to dark green or

black and is intensely bioturbated by vertical burrows ranging from 0.3 to 1.2 cm in

diameter, which weather in positive relief (Figures AI.63, and AI.64).

3.1.4 Monitor Butte and Temple Mountain Members

The Monitor Butte Member is recognized in the White Canyon, Circle Cliffs,

and Capitol Reef areas and is described in measured sections 14-27 while the Temple

Mountain Member is described in the San Rafael Swell at measured sections 1,2, and

4-13. Stewart et al. (1972) describe Monitor Butte strata in the southwestern portion

of the San Rafael Swell; however, the distinction between the Monitor Butte and the

Temple Mountain Members in this region is unclear. As a result, a lithostratigraphic

boundary separating the Monitor Butte from the overlying Temple Mountain Member

is approximated in measured sections 11-13.

In White Canyon, Capitol Reef, and the Circle Cliffs the Monitor Butte

Member is found to overlying the Moenkopi, Shinarump, or either of the soils

described in sections 3.1.1 or 3.1.3. The thickness of the Monitor Butte in these

locations varies from 93 meters at measured section 27 to zero east of measured

section 11 in the San Rafael Swell.

In general, the bottom two-thirds of the Monitor Butte consists of two overall

coarsening-upward, fluvio-lacustrine units roughly 30 meters thick. Palustrine coals,

minor delta plain paleosols, and small-scale meandering fluvial channels can be found

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near the base of the first of these sequences in the deepest portions of the paleovalley

(Figures AI.65, and AI.66); however, most often the base is characterized by thinly

laminated fine-grained deltaic facies (Figures AI.56, and AI.59), which occasionally

contain thin detrital coal beds (Figure AI.52). Above these heterolithic basal deposits

the remaining lower two-thirds of the Monitor Butte is characterized by gradually

coarsening upward beds, consisting of mudstone to medium-grained sandstone,

dominated by thin planar and current-ripple laminations (Figures AI.51, AI.56, AI.60,

and AI.69). The bulk of these deposits are interpreted as Tb and Tc turbidites typical

of delta foresets (Figure 3.1.4-1). Occasionally, slightly thicker and coarser rippled

and trough-cross bedded sandstone beds overlie the thinly bedded rocks. This facies

is interpreted as lacustrine delta topsets and lacustrine shoreface deposits.

The upper third of the Monitor Butte Member is characterized by three main

facies associated with a high-sinuosity fluvial system: 1) pedogenically modified

floodplain deposits, 2) levee and crevasse splay deposits exhibiting minor

pedogenesis, and 3) fluvial channels and point bars with variable degrees of

pedogenesis (Figures 3.1.4-2, and 3.1.4-3). These facies found in the upper portion of

the Monitor Butte are identical to those observed in the Temple Mountain Member

throughout the southern portion of the San Rafael Swell. The upper Monitor Butte

and Temple Mountain strata occupy the same stratigraphic position within the

paleovalley and, therefore, these units will be described collectively (for further

discussion see section 5.2.1).

Facies 1 consists of pedogenically modified red and purple colored

mudstones. The most common pedogenic indicators include color mottles,

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slickensides, root traces, crayfish burrows, and a lack of primary physical

sedimentary structures (Figures 3.1.4-2, and 3.1.4-3B). Vertisols, calcisols and

protosols are identified in these strata. In most locations (measured sections 5, 7, 8, 9,

11, 12, 16, 20, and 27) the amount of pedogenic carbonate increases upsection

through the Monitor Butte.

Facies 2 is characterized by 10-40 cm thick, fine- to medium-grained, current

and climbing current-rippled, red and orange sandstones interbedded with planar

bedded silty mudstones which range from 1-15 cm thick (Figure 3.1.4-3A). At Chute

Canyon in the San Rafael Swell these deposits (measured section 5, 28-39) can be

traced laterally southward to a high sinuosity fluvial channel (Figures AI.16, and

AI.17). Often these deposits are slightly pedogenically modified and contain Scoyenia

burrows, root traces, and other signs of bioturbation.

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Figure 3.1.4-1 Medium-grained sandstone preserved in the Monitor Butte Member near

Capitol Reef at MS 25 exhibiting planar laminations, current ripples, and

minor scour surfaces. Interpreted as Tb-Tc turbidite deposits associated

with a lacustrine delta front.

Figure 3.1.4-2 Reduction haloes surrounding root traces preserved in the Monitor Butte

Member in White Canyon, MS 20.

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Figure 3.1.4-3 Light red colored fine to medium grained sandstones typical of facies 2 (A)

overlain by finer-grained dark red colored vertisol typical of facies 1 (B).

Chimney Rock, Capitol Reef National Park, MS 27.

Facies 3 is made up of pink, orange, or purple colored mudstones and

sandstones exhibiting current ripples and trough-cross beds found within 3-7 meter

sized epsilon cross-stratification (ECS) sets representative of laterally-accreting point

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bar deposits. These rocks exhibit a high degree of variability both in grain size and

extent of pedogenesis. In North Wash, just north of Section 15 (Figures AI.17), at

Chute Canyon, just south of Section 5 (Figure AI.44), and at Blueberry Spring

(Figure AI.25) these rocks are moderately well sorted and range from fine- to coarse-

grained sandstones. ECS is well preserved at these locations and is the most

distinguishing characteristic of these rocks at the outcrop scale. In White Canyon, at

measured sections 17 and 20, these rocks are poorly sorted and contain over 50%

mud. At these locations, pedogenic indicators including color mottles, carbonate

nodules, burrows, and disruption of original bedding make the ECS more difficult to

see at the outcrop scale.

3.1.5 Pre-Moss Back Unconformity

In many locations (measured sections 5-7,11-13, 16-20, 22, and 27) a well-

developed carbonate-rich paleosol occupies the uppermost portion of the Monitor

Butte and Temple Mountain Members. These soils are classified as calcisols, calcic

vertisols, calcic oxisols, and calcic protosols (sensu Mack et. al, 1993). They often

contain large carbonate nodules (Figure 3.1.5-1), carbonate filled burrows (Figure

AI.15), and/or carbonate lenses. In several instances the carbonate features seem to

overprint previously developed pedogenic features characteristic of the soils

described in section 3.1.4. In these cases carbonate can be found along vertic cracks,

crayfish burrows, slickensides, and remnant bedding planes. In Capitol Reef National

park the soils found at the top of the Monitor Butte Member are characterized by

dense crayfish burrows originating at the tops of the uppermost horizon. (Figure

3.1.5-2).

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Figure 3.1.5-1 Large pedogenic carbonate nodules found in a well-developed paleosol at the

top of the Monitor Butte Member in the southwestern portion of the San Rafael

Swell, MS 11.

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Figure 3.1.5-2 Intensely bioturbated paleosol truncated by the overlying Moss Back Member

at Chimney Rock, Capitol Reef National Park, MS 27.

3.1.6 Moss Back Member

The Moss Back Member is found throughout the field area unconformably

overlying: 1) the Moenkopi Formation in the eastern portion of the San Rafael Swell

at measured section 3, 2) the Temple Mountain Member at measured sections 1, 2, 4-

10, and 3) the Monitor Butte Member in Capitol Reef, White Canyon, Circle Cliffs,

and western San Rafael Swell at measured sections 11-20, 22-27. Where present, the

Moss Back forms a distinct, resistant, brown colored cliff composed of moderately

sorted, angular, quartz sandstone conglomerate often containing numerous rounded

carbonate clasts 1-30 cm in diameter (Figures 3.1.6-1, and AI.61). The Moss Back

varies from 1-100 meters thick and can be seen nearly pinching out at several places

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near Hidden Splendor Mine in the western San Rafael Swell, and completely

pinching out in White Canyon, and east of Chimney Rock in Capitol Reef National

Park (Figures AI.53, and AI.54).

Along the south wall of Blue Notch Canyon in White Canyon, the Moss Back

forms a resistant, ledge-forming, channel shaped sandstone body incised into the

underlying Monitor Butte Member. The unit varies from over 30 meters to less than 5

meters thick and is comprised of moderately well sorted quartz-rich sandstone

deposited in large-scale poorly developed scroll bars and trough cross-beds. Aerial

photographs (Figure 5.1.5-3) support the interpretation of the Moss Back in Blue

Notch Canyon as an amalgamation of fluvial channels occupying a low-sinuosity,

steeply incised paleovalley. From above, the resistant ledge-forming sandstone can

be traced NW (253) along the southern wall and across the western end of Blue

Notch Canyon. From there the channel appears to bend northward until it goes into

the subsurface near Lake Powell. Another similar sized amalgamation of channels,

occupying a distinct lens shaped channel belt, can be seen in outcrop, oriented NNW

(349), further south near measured section 20.

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Figure 3.1.6-1 Aerial photograph of Blue Notch Canyon showing extent of Moss Back

Member (basemap from www.maps.google.com).

North of Blue Notch Canyon at Copper Point, the Moss Back is comprised of

less than 10 meters of coarse graveley conglomerate containing numerous carbonate

nodule and mud rip-up clasts (Figure 3.1.6-1). In North Wash, the Moss Back once

again consists of amalgamated fluvial channels made up of large-scale, trough cross-

bedded and moderately well-sorted quartz sandstone; however, the thickness of the

unit is less variable (12-20 meters) than further south and the overall form of the

sandbody is tabular rather than channel (lens) shaped.

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Figure 3.1.6-2 Thin deposit of conglomerate comprising the Moss Back Member at Copper

Point, MS 16 (A) truncating a well developed paleosol exhibiting densely spaced

carbonate nodules (B). Overlain by dark red, pedogenically modified high-

sinuosity fluvial sandstones of the Petrified Forest Member (C).

In the southwestern portion of the San Rafael Swell, near Hidden Splendor

Mine, the Moss Back sandbody is once again channel shaped and ranges from 6-25

meters thick; however, the unit quickly becomes tabular to the east. From Bell

Canyon to Temple Mountain (measured sections 4-10), the Moss Back is made up of

amalgamated moderate-sinuosity fluvial channels. North of Temple Mountain at

measured sections 2, 3 and surrounding section 1, the Moss Back increases in

thickness up to 90 meters and completely truncates all of the Temple Mountain

Member in some locations. Unlike exposures in White Canyon, the Moss Back

structurally dips over 45 degrees in this region providing only 2 dimensional

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exposures. Deposition of the lowermost portion of the Moss Back in the Three

Fingers Canyon region was laterally constrained by deeply incised paleotopography.

Photopans and paleocurrent measurements taken at Spotted Wolf Canyon several

miles to the north suggest the upper portion of the Moss Back in the eastern portion of

the Swell was dominated by lateral fluvial channel amalgamation similar to the Moss

Back exposed further south.

3.1.7 Petrified Forest Member

The Petrified Forest Member conformably overlies the sandstone

conglomerate of the Moss Back Member. Where the Moss Back is not present

(sections 23, and 32) the Petrified Forest Member overlies a distinctive, dark red

colored paleosol at the top of the Monitor Butte Member.

In the San Rafael Swell the Petrified Forest Member is characterized by a

color change from white or light brown to dark red and by a change in fluvial style.

While the Moss Back contains trough cross-beds, ripples and massively bedded units

deposited in distinct channel forms and lateral bar deposits, the base of the Petrified

Forest consists of a distinct red colored, well sorted trough and recumbent cross-

bedded tabular sandstone bodies (Figure 3.1.7-1) which overlap the amalgamated

Moss Back channels and correlative interfluves. These sandstones correlate laterally

to dark red pedogenically modified mudstones at Spotted Wolf Canyon in the San

Rafael Swell, and throughout most of White Canyon and Capitol Reef.

In White Canyon and Capitol Reef National Park, the base of the Petrified

Forest Member is recognized by its distinct red color; however, these strata tend to be

much finer-grained overall than those preserved in the San Rafael Swell. Here the

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strata consist of pedogenically modified mudstones and fine sandstones that exhibit

popcorn (badlands style) weathering. These deposits are poorly sorted and show

remnant planar bedding, trough cross-bedding, current ripples and climbing ripples,

and 2 meter high lateral accretion sets (Figure 3.1.7-2).

Figure 3.1.7-1 Well sorted trough cross-bedded sandstones found near the base of the Petrified

Forest Member near MS 5 in the San Rafael Swell.

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Figure 3.1.7-2 Fine-grained, red colored paleosol exhibiting relict lateral accretion sets found

at the base of the Petrified Forest Member at MS 18.

3.2 Paleocurrent Data

Rose diagrams and statistical analyses of paleocurrent measurements are shown

in Figure 3.2-1 and Appendix II. The average paleocurrent vector for the lower

portion of the Chinle, 330, is consistent with the findings of Stewart et al. (1972) and

others. The Shinarump Member averaged 345, the Monitor Butte and Temple

Mountain Members averaged 357, and the Moss Back and lower Petrified Forest

Members averaged 313.

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Figure 3.2-1 Spatial and stratigraphic distribution of paleocurrent directions measured in

the lower portion of the Chinle Formation in south central Utah.

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3.3 Petrographic Data

The results of the point count analysis can be seen in Figure 3.2-1 and in

Appendix III where sandstones are plotted on QFL ternary diagrams and classified on

the basis of composition (after Pettijohn, 1975).

The Shinarump, Moss Back, and Lower Petrified Forest Members contain

arkosic, subarkosic, and quartz arenites. The Temple Mountain and Monitor Butte

Members contain arkosic, and subarkosic arenites and sublitharenites. Overall, the

maturity of the sandstones decreases and becomes more variable up section.

Spatial trends in sandstone composition are also observed. The Shinarump,

Monitor Butte, and Temple Mountain Members become less mature downstream

from White Canyon to the Circle Cliffs to Capitol Reef. The Moss Back Member

exhibits more spatial variability; however, sandstone compositions become slightly

more mature downstream.

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Figure 3.3-1 Relative abundances of total quartz, total feldspar, and total lithic clasts as

determined by point-count analyses of sandstones from each of the three

depositional sequences outlined in Chapter 4.

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CHAPTER IV

SEQUENCE STRATIGRAPHY

4.1 Introduction to Sequence Stratigraphy

Sequence stratigraphy was developed in the late 1970s promoting the

widespread analysis of sedimentary deposits within a time-based stratigraphic

framework. Whereas traditional lithostratigraphic methods correlate groups of rocks

based upon similar physical characteristics, sequence stratigraphy acknowledges

Walthers law of facies by grouping genetically related packages of rocks

(depositional sequences, systems tracts, and parasequences) which are separated

by chronostratigraphically-significant surfaces (flooding surfaces, and sequence

boundaries) (Figure 4.1-1). Once identified, the geometries and stacking patterns

(stratal architecture) of these packages can be used to interpret a depositional history

based upon fluctuations in accommodation through time. Models developed as a

result of the application of this methodology have proven useful as a predictive tools

for hydrocarbon exploration and have helped to constrain the timing and relative

importance of various forcing mechanisms such as climate, tectonics, and

eustasy/limnostasy (Keighley et al., 2003).

depositional sequence (DS): A stratigraphic unit bounded at its top and base by unconformities or correlative conformities.

flooding surface (FS): A surface separating older from younger rock marked by deeper-water strata resting on shallower-water

strata.

highstand systems tract (HST): The portion of a depositional sequence characterized by aggradation and progradation which

is bounded below by the maximum flooding surface and above by the overlying sequence boundary.

lowstand systems tract (LST): The portion of a depositional sequence that is bounded below by a sequence boundary and

above by a transgressive surface.

maximum flooding surface (MFS): A flooding surface used to separate transgressive and highstand systems tracts identified bya change from retrogradation to aggradation/progradation. Marks the time when the shoreline is at its maximum landward

position, thus subsequent strata downlap onto this surface.

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parasequence: A genetically related package of rocks that is relatively conformable and is bounded above and below by

flooding surfaces.

sequence boundary (SB): A chrono-stratigraphically significant surface characterized by erosional truncation and basinward

shift of facies. Overlying strata onlap this surface.

transgressive surface (TS): The first major flooding surface used to separates late LST from TST. Subsequent stratal onlapmarks the trend of a shoreward moving shoreline.

transgressive systems tract (TST): The portion of a depositional sequence bounded below by the transgressive surface andabove by the maximum flooding surface marked by landward shift in facies.

Walthers law: Different sediment types (facies) accumulate beside each other at the same time, therefore vertical lithologic

changes in the rock record reflect lateral shifts of facies through time (assuming continuous deposition).

Figure 4.1-1 Glossary of sequence stratigraphic terminology introduced in section 4.1 (Fritz

and Moore 1988, Nichols, 1999, Posamentier and Allen, 1999)

4.2 Application to Continental Strata

Fluvial Deposits. Sequence stratigraphy has been most successfully applied to

nearshore marine sedimentary deposits affected by eustasy (Posamentier and Allen,

1999), where the affects of variations in base level and accommodation through time

can be readily recognized in the rock record. While traditional sequence stratigraphic

methods have been successfully applied to fluvial deposits isolated from the effects of

base level, interpretations become difficult due to the uniquely dynamic nature of

fluvial systems. Unlike nearshore deposits where accommodation (depositional

potential) is dependant upon relative sea level, accommodation in fluvial systems is

either gained or lost by positive and negative shifts in the equilibrium profile of a

stream system, respectfully (Shanley and McCabe, 1994). Changes in basin length,

sediment supply, sediment size, discharge, and base level drive shifts in the fluvial

profile, which ultimately lead to either aggradation or incision of the fluvial surface

(Figure 4.2-1) (Lane, 1955, Allen, 1990, Quirh, 1996, Blum and Tornqvist, 2000).

Overall, fluvial systems are more sensitive to upstream forces (especially sediment

supply) than are marine deposits (Shanley and McCabe, 1994).

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Figure 4.2-1. Model depicting the controls on aggradation and degradation in alluvial

systems. If sediment supply out-weighs stream power the system will aggrade.

Degradation will occur if the system is sediment starved in comparison to the

stream power (after Lane, 1955).

Marine facies assemblages are generally differentiated based upon water

depth indicators and serve as a proxy for accommodation in marine systems; in

contrast, fluvial systems exhibit complicated and highly variable facies assemblages

that may not be directly related to accommodation. While marine facies vary along

dip and are dependent upon water depth, continental depositional environments vary

along strike primarily due to variations in topography. As a result, the application of

Walthers Law becomes difficult to apply because shifts in facies cannot be used as a

proxy for changes in accommodation.

Despite these inherent difficulties several workers have proposed sequence

stratigraphic methods for fluvial systems that use channel stacking patterns and

degree of pedogenesis as proxies for sediment accumulation rates (accommodation)

(Wright and Marriott, 1993, Shanley and McCabe, 1993, 1994). In general these

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studies suggest dense channel stacking and increased pedogenesis as evidence of low

accommodation while finer grained deposits with higher overbank/channel ratios and

lower degree of pedogenesis suggest high accumulation rates resulting from increased

accommodation.

Lacustrine Deposits. Accommodation is more easily constrained for

lacustrine deposits where many facies are water depth dependent. Water depth is a

function of basin morphometry (tectonically controlled) and sediment + water supply

(climate controlled). These two factors serve as the basis for the model presented by

Carrol and Bohacs (1999) (Figure 4.2-2) which categorizes lacustrine basins into

underfilled, balance- filled, and overfilled basins based upon lacustrine facies

assemblages.

Figure 4.2-2 Lacustrine classification scheme proposed by Carroll and Bohacs (1999) based

upon the relative influence of tectonic and climatic controls.

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Sequence stratigraphic methodology has successfully been used to construct

an accommodation history for siliciclastic fluvio-deltaic lacustrine deposits by

considering limnostatic base level fluctuations (e.g. Lemons and Chan, 1999). As in

nearshore deposits, flooding surfaces can be identified and correlated a short distance

landward by examining paleohydrologic indicators preserved in paleosols formed in

adjacent lowlands. However, several distinct differences between limnostatically and

eustatically controlled systems must be considered. First, unlike rivers that flow into

oceans, the discharge of inflowing streams partially controls lake level. Secondly,

lakes are relatively short-lived features and lake level fluctuations can occur on much

shorter time scales (Keighley et al., 2003). An example of these differences can be

seen in underfilled lacustrine basins where short term transgressions during storm

events create fining-upward parasequences (Smoot and Lowenstein, 1991). Overfilled

lake basins also differ from marine basins because tectonic forces, acting to either

raise or lower the sill, affect limnostasy (Carroll and Bohacs, 1999, Keighley et al.,

2003).

4.3 Discussion of Unconformity Paleosols

The well-developed paleosols described in sections 3.1.1, 3.1.3, and 3.1.5

were formed along the margin of the paleovalley (interfluves) during long periods of

local stability or degradation (McCarthy and Plint, 1998, Demko et al., 2004). The

timing and duration of paleosol development is difficult to constrain in these soils,

however they can be used to help identify and correlate several

chronostratigraphically-significant surfaces (sequence boundaries) within the lower

portion of the Chinle.

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Determining the relative timing of interfluve paleosol development and

adjacent deposition is difficult because the two may occur contemporaneously. For

example, while deposition of the Shinarump was occurring in the center of the

paleovalley near Capitol Reef National Park, interfluve pedogenesis, and perhaps

minimal erosion, continued across the remaining portion of the landscape. This

spatial variability makes it difficult to constrain when each paleosol started to form.

For example, the paleosol found at the base of the Chinle at MS 4 may have started to

form as early as the Middle Triassic after merely 10 meters of Moenkopi was incised

or as late as sometime during the deposition of the lower portion of the Monitor Butte

at the base of the paleovalley (early DS 2).

Soil thickness (depth of weathering) can be used to help make a first order

approximation of the age of an interfluve paleosol by assuming pedogenesis occurred

from the top down. However, since these paleosols spend much of their time forming

on a degrading landscape, and correlate laterally to eroded strata, it is likely that much

of the soil was removed prior to burial.

These problems are addressed when using these palosols as

chronostratigraphic indicators by making several key assumptions about when and

where pedogenesis occurs:

1) Deposition does not occur on the interfluve, the deposits neatly onlap

the paleovalley from the bottom up.

2) Pedogenesis may occur continuously, both during aggradational and

degradational time periods, at all locations above where the strata is

currently onlapping the paleovalley margin.

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These assumptions serve to temporally isolate the deposition of parent

material from subsequent pedogenesis creating an accurate time-depositional

framework for the Chinle Formation in which the tops of the paleosols are used to

delineate depositional time periods separated by periods of landscape stability or

degradation. The implicit time-pedogenic framework constrains paleosol

development to a period ofpotentialpedogenesis bounded by a maximum age (after

deposition of the original parent material) and a minimum age (just before the soil is

onlapped by subsequent deposition).

4.4 Sequence Stratigraphic Interpretation

Section 4.4 applies sequence stratigraphic methodology to the lower portion

of the Chinle Formation. Three sequence boundaries (SB 1-3) are identified and

correlated across the field area serving to divide the strata into 3 depositional

sequences (DS 1-3) (Figures 4.4-1 and 4.4-2 ). These three sequences closely

resemble the established lithomember classification for the lower portion of the

Chinle, however several key differences exist. The most notable difference occurs at

the Tr-3 unconformity where the sequence boundary is interpreted at the top of well

developed interfluve paleosol whereas the lithologic contact has been defined at the

base of the soil (see measured sections 1,4, 6-7, 9, 10, 13, 22).

Further identification of lacustrine flooding surfaces, their correlative

surfaces, and paleohydrologic indicators preserved in paleosols, provides insights to

the accommodation history (changes in relative lake level) of the basin. These data, in

addition to observed stratal architecture (both at outcrop scale and in interpreted

cross-sections), permits the division of DS 2 into three secondary (higher order)

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Sequence Stratigraphy of Lacustrine