T-4010

STRATIGRAPHY OF THE DEVONIAN GUILMETTE FORMATION, PAHRANAGAT RANGE, LINCOLN COUNTY, NEVADA

by Jane E. Estes

b o a t e r ProQuest Number: 10783700

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ProQuest LLC. 789 East Eisenhower Parkway P.O. Box 1346 Ann Arbor, Ml 48106- 1346 A thesis submitted to the Faculty and the Board of Trustees of the Colorado School of Mines in partial fulfillment of the requirements for the degree of Master of Science

(Geology).

Golden, Colorado

Date Q j u u i m z -

Signed: Jafle E. Estes

Approved: Dr.yTdhn E. Warme £sis Advisor

Golden, Colorado Date hpfx'l V

Dr. Gregory S. Holden Acting Head, Department of Geology and Geological Engineering T-4010

ABSTRACT

The Devonian Guilmette Formation is well exposed in the Pahranagat Range,

Lincoln County, Nevada. The purpose of this study is to describe and interpret deposi- tional sequences within the Guilmette in this area.

The Guilmette Formation was subdivided into three Members based on its lithology and depositional processes. The Lower Member is composed of interbedded limestone and dolomite deposited as shallowing-upward cycles in a shallow platform environment. The

Middle Member is separated from the Lower and Upper Members by sharp, erosional contacts. This enigmatic Member exhibits characteristics of submarine debris flows and was deposited in relatively deep water. The Upper Member is composed of interbedded limestone, dolomite, and sandstone deposited in subtidal to supratidal environments.

These units appear to be somewhat cyclic, although not as well defined as the cycles of the

Lower Member.

In terms of sequence stratigraphy, the Lower Member is interpreted as the transgressive systems-tract due to its overall deepening-upward nature. The Middle

Member is thought to represent a highstand-systems tract because the accommodation space necessary for this deposit to accumulate requires a very high stand of relative sea level. The top of the Middle Member, which is thought to represent a relative sea level fall, is interpreted as a type 2 sequence boundary. The Upper Member appears to be regressive overall and is interpreted as the shelf margin wedge. Computer-generated Fischer plots were used as an objective means of supporting this interpretation.

iii T-4010

Synthetic gamma ray curves constructed for the Guilmette in the study area indicate regional trends that may be useful for exploration. In general, the Lower Member exhibits high gamma values and the Middle Member exhibits low gamma values. Additionally, unconformity at the top of the Middle Member is marked by a spike on the gamma curve that may serve as a regional marker.

iv T-4010

TABLE OF CONTENTS

page ABSTRACT ...... iii

LIST OF FIGURES ...... vii

LIST OF PLATES ...... ix

ACKNOWLEDGEMENTS ...... x

Chapter 1. IN T R O D U C T IO N ...... 1

Purpose and Scope ...... 1 Location of Study A rea ...... 1 Methods of Investigation ...... 2 Explanation of Plates ...... 4 Regional Setting ...... 5

2. PREVIOUS W ORK ...... 13

Stratigraphic Studies ...... 13

Paleoecological Studies...... 17 3. LOWER MEMBER OF THE GUILMETTE FORMATION 19

Introduction ...... 19 L ith o lo g y ...... 19 Petrography ...... 22 Environment of Deposition ...... 32

4. MIDDLE MEMBER OF THE GUILMETTE FORMATION ... 36

Introduction ...... 36 Lithology ...... 36 Petrography ...... 43 Environment of Deposition ...... 49 v T-4010

5. UPPER MEMBER OF THE GUILMETTE FORMATION 52

Introduction ...... 52 Lithology ...... 52 Petrography ...... 56 Environment of Deposition ...... 65

6. SEQUENCE STRATIGRAPHIC FRAMEWORK OF THE GUILMETTE FORMATION ...... 71

Introduction ...... 71 Key Definitions...... 71 Fischer P lots ...... 72 Interpretation ...... 76 7. PETROLEUM POTENTIAL OF THE GUILMETTE FORMATION ...... 80

Introduction ...... 80 Exploration ...... 80 Reservoir ...... 83

8. SUMMARY AND CONCLUSIONS...... 84

REFERENCES CITED ...... 87

vi T-4010

LIST OF FIGURES

page

1.1 Location map ...... 3

1.2 Devonian paleogeographic and paleotectonic map ...... 7

1.3 Geologic map of study area ...... 12

2.1 Outcrop photo showing three Members of the Guilmette ...... 18

3.1 Basal contact of the Guilmette ...... 20

3.2 Dolomite of the Lower Member...... 22

3.3 Lime oncoid packstone bed ...... 23

3.4 Photomicrograph of SMF 9 in Lower Member ...... 25

3.5 Photomicrograph of SMF 16 in Lower Member ...... 26

3.6 Photomicrograph of SMF 19 in Lower Member ...... 27

3.7 Photomicrograph of SMF 20 in Lower Member ...... 28

3.8 Photomicrograph of SMF 23 in Lower Member ...... 30

3.9 Photomicrograph of diagenetic phases in Lower Member ...... 31

3.10 Idealized cycle of the Lower M ember...... 33

4.1 Basal contact of the Middle Member...... 37

4.2 Typical clast assemblage of the Middle Member...... 38

4.3 Large dolomite clasts within the Middle Member ...... 40

4.4 Lens-shaped bed of stromatoporoids within the Middle Member ...... 41

4.5 Stromatoporoid biostrome within the Middle M ember ...... 42

vii T-4010

4.6 Photomicrograph of SMF 4 in Middle Member ...... 44

4.7 Photomicrograph of SMF 5 in Middle Member ...... 46

4.8 Photomicrograph of SMF 20 in Middle M ember ...... 47

4.9 Outcrop photo of Middle Member showing beds and bedsets...... 50

5.1 Bioturbated lime wackestone of the Upper Member...... 54

5.2 Fossiliferous lime packstone of the Upper Member ...... 55

5.3 Trough cross-bedded sandstone of the Upper Member ...... 57

5.4 Algal-bound sandstone of the Upper Member ...... 58

5.5 Eolian sandstone of the Upper Member ...... 59

5.6 Photomicrograph of SMF 16 in Upper Member ...... 61

5.7 Photomicrograph of SMF 19 in Upper Member ...... , ...... 62

5.8 Photomicrograph of SMF 20 in Upper Member ...... 63

5.9 Photomicrograph of SMF 23 in Upper Member ...... 64

5 .1 0 Photomicrograph of SMF 25 in Upper M em ber ...... 66

5.11 Common facies associations within the Upper M ember ...... 68

6.1 Fischer plot example ...... 74

6 .2 Interpreted age of the Guilmette Formation ...... 75

6 .3 Fischer plot of the Guilmette Formation in thePahranagat R a n g e ...... 77

6 .4 Relative sea level curve for the Guilmette in the Pahranagat R an g e ...... 78

7.1 Petroleum potential map of Nevada ...... 82

viii T-4010

LIST OF PLATES

PLATE 1 ...... in pocket

PLATE 2 ...... in pocket

ix T-4010

ACKNOWLEDGEMENTS

I would first like to thank my advisor, Dr. John Warme, and my committee members, Dr. Mark Longman and Dr. Trobe Grose, for their guidance and insight.

Unocal Corporation provided the funding for this project, and I thank Norman Kent of

Unocal for his time in the field and his interest in the project. I am deeply indebted to Brian

Ackman and Alan Chamberlain for their assistance in the field and otherwise. I could not have done it without their help.

This thesis has greatly benefited from discussions with numerous individuals, including Rick Geesaman, Jim Wilson, Peter Isaacson, David Valasek, Mark Hanson,

Mark Sonnenfeld, Keith Shanley, and Lyn Canter. However, the interpretations presented herein are solely the responsibility of the writer. The Fischer plot computer program was provided by Charles Kerans through Coyote Geologic Services.

I am very grateful to all of my family and friends for their encouragement, and I am especially thankful to my mother Mary Wyche Estes for supporting me both morally and financially during my academic career. This thesis is dedicated to the memory of my father

Glen Estes.

Finally, I would like to thank Todd Jackson for helping me to maintain my perspective through the past three years. This project would not have been possible without his infinite patience, encouragement, and humor.

x T-4010 1

CHAPTER 1 INTRODUCTION

Purpose and Scope

The Guilmette Formation is an outstanding example of a Devonian carbonate depositional system. While numerous workers have focused on the gross stratigraphy of the

Guilmette, few have documented the complex facies changes within the unit. In addition,

Devonian carbonates in the Great Basin region are important economically because they provide reservoirs for significant quantities of oil. At Grant Canyon Field, in Nye County,

Nevada, the Guilmette Formation has cumulatively produced over 13.8 million barrels of oil from two wells since its discovery in 1983. Average daily production in this field is over

7,000 BOPD (Petzet, 1991).

Well exposed outcrops of the Guilmette Formation in southeastern Nevada provide an excellent opportunity to study this unusual carbonate system. The purpose of this reconnais sance study is to describe and interpret depositional sequences within the Guilmette in the Pahranagat Range, Lincoln County, Nevada. The primary objectives of this study are:

1) to describe the facies tracts and depositional sequences within the Guilmette, 2) to interpret the Guilmette using sequence stratigraphic analysis, and 3) to use this information to predict potential hydrocarbon reservoirs in the Guilmette.

Location of Study Area

The study area is on the west side of the Pahranagat Range in Lincoln County,

Nevada and is reached by driving west from Crystal Springs approximately 15 miles on T-4010 2

Nevada 375 and then by driving 1 mile south and east up a drainage on a four wheel drive trail that intersects the paved road approximately 1 mile west of Hancock Summit. This area is approximately 60 miles south-southeast of Grant Canyon Field (Figure 1.1). The

Pahranagat Range probably contains the thickest and most complete Devonian section in the

Great Basin region (Tschanz and Pampeyan, 1970), and was chosen for this study because of the excellent outcrop exposures of the Guilmette and the relative accessibility there. One stratigraphic section (Section 1, Plate 1) was measured in Sections 20,21, and 29, Township

6 South, Range 59 East, and another stratigraphic section (Section 2, Plate 2) was measured in Section 5, Township 7 South, Range 59 East. Section 1 comprises the entire Guilmette interval, with no repetition or omission due to faulting, and consists of approximately 2300 feet of Guilmette. A complete section was not measured at Section 2, however, due to inaccessibility.

Methods of Investigation

Both stratigraphic sections in this study were measured during late May through early June of 1990, using the Jacob staff method. Carbonate rocks were described using

Dunham’s (1962) classification and clastic rocks were described using the classification of

Dott modified by Pettijohn, et al. (1987). The gamma values derived from a scintillation counter were used to construct a synthetic gamma ray log for each measured section. In both measured sections, chip samples were collected every 20 feet, and hand-sized samples were collected every 100 feet and at major stratigraphic contacts. Those chip samples that were thought to best represent lithology and depositional environment were cut, stained with

Alizarin Red to determine calcite content, and then made into thin sections. A total of 64 thin sections were examined microscopically to verify field descriptions of lithology, were T-4010 3

■WhftBifine County*

Grant Cannon* Bacon Flat i

NEVADA

Nye County^ Lincoln County

Crystal Springs'

0 10 20 3 0 40 50 ______1 I______I______I______I______I Miles FIGURE 1.1: Map showing location of study area in relation to major ranges and Railroad Valley oil fields. See Figure 1.3 for a detailed map of the study area. T-4010 4 further analyzed for primary texture, diagenetic history, and porosity, and then were classified using Wilson’s (1975) Standard Microfacies. These microfacies were then used to augment the interpretation of the facies and depositional environments of the Guilmette.

Hand samples were cut and polished and used to identify macrofossils and depositional textures. The facies and depositional cycles defined within the Guilmette were used to interpret the formation in a sequence stratigraphic framework, which was then used to construct a relative sea level curve for the region. A weakness of this sea level curve is the lack of biostratigraphic control that could place it into a global time framework.

Explanation of Plates

The two stratigraphic sections measured for this study are included as plates located in the pocket behind the text. The reader is encouraged to refer to these plates while reading the text. Section 1 is presented in Plate 1 and Section 2 is presented in Plate 2. Both sections were drafted at the scale of 1 inch equals 20 feet in order to present a detailed representation of the rocks. The plates were reproduced at the scale of 1 inch equals 40 feet to be more convenient for the reader. A synthetic gamma ray log constructed from field measurements was plotted along the left side of each measured section. Interpreted cycle tops, represented by small arrowheads, were plotted immediately to the right of the lithologic column. The large arrowheads represent major stratigraphic contacts. Fossils observed in the field are plotted in columns to the right of the lithologic column. In these columns, the thick line represents abundant fossils, the thin line represents fossils present, and the dashed line represents rare fossils. Thin section locations and their sample numbers were plotted to the right of the fossil columns. Field descriptions of each measured section were placed on the far right of each plate. Descriptions were subdivided to correspond to changes in lithology, T-4010 5 and were verified by thin sections and samples, where available.

The traverse for Section 1, as well as formation contacts and strike and dip symbols, were plotted on the index map of Plate 1. Section 1, informally called “Arrowhead Ridge” by the field party, was started at the base of the Fox Mountain Member of the Simonson

Dolomite in order to better describe the relationship of the Guilmette to the Simonson.

Approximately 2300 feet of Guilmette was measured at Arrowhead Ridge. The top of

Section 1 was placed at the contact of the overlying Pilot Shale, which forms a strike valley between the more resistant Guilmette and overlying Mississippian Joana Limestone.

The traverse for Section 2, as well as formation contacts and strike and dip symbols, were plotted on the index map of Plate 2. Section 2, informally called “Devil’s Doorstep” by the field party, was started in the Simonson Dolomite approximately 40 feet below the contact with the Fox Mountain Member. The top of this section was placed approximately

35 feet above the top of the Middle Member of the Guilmette because of inaccessibility.

Regional Setting

The following discussion was compiled from various authors and is intended to be a summary.

Paleogeogranhv: From Late Precambrian through Late Devonian time, the western

United States was the site of thick sediment accumulation in a passive margin setting

(Burchfiel and Davis, 1972; Dickinson, 1977). The Guilmette was deposited on this shelf as shallow-water cyclic carbonates (Cook, 1988; Sandberg, etal., 1988). The study area was located on the shelf at an approximate paleolatitude of 15° S, which is within the arid climate belt (Witzke and Heckel, 1988). To the west of this miogeosyncline was an eugeosyncline, composed of the inner arc basin in eastern Washington, eastern Oregon, western Nevada, and T-4010 6 northern California, and the island arc chain, near the present-day western coast of the United

States. The cratonic platform was located to the east of the study area, in central and eastern

Montana, Wyoming, eastern Utah, and Arizona (Figure 1.2; Poole, et al., 1967; Poole, et al.,

1977).

Paleotectonics: The tectonic history of the Great Basin region is extremely complex; however, it may be subdivided into three general phases of evolution, as outlined by Burchfiel and Davis (1972), Dickinson (1977), Cook (1988), andLevy and Christie-Blick

(1989): 1) Late Precambrian through Devonian continental extension and development of a passive continental margin, 2) Late Devonian through Eocene crustal shortening, and 3)

Late Eocene through Recent crustal extension. A Late Precambrian rifting event initiated the development of the proto-Pacific Ocean and a north-trending passive continental margin in western North America. The edge of the continent at that time is thought to be represented by the 87Sr/86Sr = 0.706 contour line, located approximately at present longitude 117° W

(Cook, 1988; Sandberg, et al., 1988; Levy and Christie-Blick, 1989). As previously discussed, the Cordilleran miogeosyncline developed on this passive margin. A thick sequence (up to 16,000 ft) of Cambrian through Devonian carbonates was deposited on this shelf, while deep-water sediments were deposited in the ocean basin to the west (Poole, et al., 1967; Burchfiel and Da vis, 1972; Dickinson, 1977; Poole, etal., 1977; Cook, 1988; Levy and Christie-Blick; 1989).

The second phase of Great Basin evolution occurred during the Late Devonian through Eocene, and consists of four crustal-shortening orogenies: 1) the Antler Orogeny

(Late Devonian-Early Mississippian), 2) the Sonoma Orogeny (Permian-Triassic), 3) the

Sevier Orogeny (Cretaceous), and 4) the Laramide Orogeny (Late Cretaceous-Eocene)

(Armstrong, 1968; Burchfiel and Davis, 1972; Dickinson, 1977; Cook, 1988; Levy and T-4010 7

127° 124“ 121“ 118“ 115“ 112“

48'

MONTANA 48' WASHINGTON

45' IDAHO

45 '

OREGON

4 2'

4 2'

UTAH

CALIFORNIA NEVADA Xj 39'

39'

ay _ A rea I

36 '

36 ‘ ARIZONA

33 ‘ SCALE 100 150 200

33'

124“ 118“ 115“ 112“ FIGURE 1.2: Devonian paleogeographic and paleotectonic map of the western United States showing major post-Devonian faults and location of study area (modified from Poole, et al., 1977). T-4010 8

Christie-Blick, 1989). The Antler Orogeny marks the change from passive to collisional continental margin (Goebel, 1991), and is manifested in the Roberts Mountain Allochthon in north central Nevada, which is composed of Cambrian through Devonian chert, argillite, greenstone, sandstone, and quartzite of the Cordilleran eugeosyncline that have been thrust eastward approximately 50 miles over the adjacent miogeosynclinal rocks (Poole, et al.,

1967; Burchfiel and Davis, 1972; Poole and Sandberg, 1977; Nilsen and Stewart, 1980;

Cook, 1988). The northeast-trending Antler Orogenic Belt appears to parallel the eastern margin of the eugeosyncline (Poole, et al., 1967). Several models for the origin of this orogeny have been proposed (Nilsen and Stewart, 1980); however, the two most accepted models are 1) back-arc thrusting associated with an east-dipping subduction zone, and 2) arc- continent collision associated with a west-dipping subduction zone (Burchfiel and Davis,

1972; Dickinson, 1977; Poole and Sandberg, 1977; Cook, 1988; Goebel, 1991).

Most recently, Burchfiel andRoyden (1991) suggested that the Antler Orogeny may be similar to Mediterranean-type orogenies “in which thrusting occurred behind a zone of trench retreat, while a region of extension developed contemporaneously within the hanging wall of the subduction systems.” In this model the thrust belt, composed of accreted rocks, developed east of the subduction zone, and trench-retreat caused this thrust belt allochthon to migrate eastward. Flexural loading of the allochthon produced an asymmetrical foreland basin to the east of the allochthon (Poole and Sandberg, 1977; Cook, 1988; Goebel, 1991).

To the east of this foreland basin a forebulge developed with a shallow back-bulge basin developing further east (Goebel, 1991). In the Late Devonian bedded cherts andhemipelagic claystones of the Pinecone Sequence and the Woodman Formation were deposited in the foreland basin, while the Pilot Shale and West Range Limestone were deposited in the back- bulge basin. During the Early Mississippian, the forebulge migrated eastward into Utah, and T-4010 9 eastern Nevada became part of the foreland basin, which was eventually filled completely by clastic sediments derived from the Antler Highland (Burchfiel andDavis, 1972; Poole and

Sandberg, 1977; Nilsen and Stewart, 1980; Cook, 1988; Levy and Christie-Blick, 1989;

Goebel, 1991).

The Permian-Triassic Sonoma Orogeny thrust the Golconda Allochthon eastward over the western edge of the rocks deformed by the Antler Orogeny (Burchfiel and Davis,

1972; Dickinson, 1977). Rocks in the Golconda Allochthon consist of oceanic assemblages and volcanic terranes that were originally an island arc complex that was welded onto the continent during this orogeny (Dickinson, 1977).

The Cretaceous Sevier Orogeny formed a northeast-trending belt of folds and thrusts that generally coincide with the eastern limit of the Cordilleran miogeosyncline (Figure 1.2).

This belt extends 500 miles from southeast Nevada to Montana, and represents 40-60 miles of shortening (Armstrong, 1968). The highland that resulted from this orogeny shed clastic sediments into the adjacent Rocky Mountain geosyncline to the east (Armstrong, 1968). The

Laramide Orogeny began almost immediately after the Sevier Orogeny ended in the Late

Cretaceous, and continued into the Eocene (Armstrong, 1968). Both orogenies resulted from the evolution of the western continental margin into an Andean-type subduction regime

(Cook, 1988). While the Laramide Orogeny is generally associated with the uplift of the

Rocky Mountains to the east, there is also evidence of Laramide structures in southeastern

Nevada (Tschanz and Pampeyan, 1970).

The third phase of the evolution of the Great Basin began in the Late Eocene and continues today (Levy and Christie-Blick, 1989). It is marked by extensional tectonics

(Wernicke, 1981) related to changes in plate interactions (Burchfiel and Davis, 1972). Eaton

(1979) described three types of extension that occurred in the region after the Laramide T-4010 10

Orogeny: 1) convergence-related intra-arc extension, 2) convergence-related back-arc extension, and 3) extension related to the cessation of plate convergence and the initiation of lateral slip of the Pacific plate past the North American plate. This extension resulted in the normal and strike-slip faults that formed the alternating horst and graben structures characteristic of the Basin and Range region (Stewart, 1971; Burchfiel and Davis, 1972;

Wernicke, 1981; Cook, 1988; Levy and Christie-Blick, 1989). This most recent tectonism has overprinted all previous orogenic episodes. Estimates of total extension range from 10-

35 percent to 100 percent depending upon the region and angle of faulting (Cook, 1988).

Hamilton (1987) estimated that the width of the province has doubled since the Oligocene.

Structural Setting of the Study Area:The study area was apparently structurally unaffected by the Antler Orogenic belt to the west. The Sevier Orogenic Front is located to the east of the study area (Figure 1.2). The Pahranagat Range formed as a result of extensional faulting during the Late Miocene (Jayko, 1988). The following discussion on the structure of this Range is summarized from Tschanz and Pampeyan (1970). The

Pahranagat Range strikes generally north-south, is approximately 36 miles long, and is comprised of three structural blocks. The westernmost block is an east-dipping homoclinal sequence of Paleozoic rocks, ranging from Cambrian to Lower Mississippian in age.

Immediately east of this block is a pre-Miocene thrust plate that places Cambrian over

Mississippian strata. This thrust plate is thought to correlate to the thrust plate in the Golden

Gate Range to the north. Tschanz and Pampeyan (1970) interpret this thrust as a Laramide feature, but Chamberlain (in prep.) interprets it as Mesozoic in age. The easternmost structural block is the East Pahranagat Range and is comprised of an overturned syncline referred to as the Pahranagat Fold. This fold affects formations ranging from Devonian to

Pennsylvanian in age, and can also be traced to the Golden Gate Range. The stratigraphic T-4010 11 sections for this study were measured in relatively undeformed Devonian rocks of the western structural block. This block contains at least 10 northeast-striking Tertiary faults probably related to Basin and Range extensional tectonics (Figure 1.3). The southern end of the Range is terminated by three northeast-striking left lateral strike-slip faults referred to as the Pahranagat Shear Zone, which is thought to be a post-Laramide reactivated

Laramide shear system that formed as a means of stress relief during the displacement of the previously mentioned thrusts (Tschanz and Pampeyan, 1970). Small-magnitude earth­ quakes within the shear zone suggest that fault activity continued through the Quaternary

(Jayko, 1988). T-4010 12

EXPLANATION Q al« Quaternary alluvium Nhb - Neogene Hells Bells Canyon Formation Mj - Mississippian Joana Lim estone DMp - Dev-Miss Pilot Shale

- Devonian Guilmette Form ation

D si« Devonian Simonson Dolomite Dse - Devonian Sevy Dolomite SI « Silurian Laketown Dolomite Oes - Ordovician Ely Springs Dolomite O e« Eureka Quartzite Op ■ Ordovician Pogonip Group (und.) Cdv - Cambrian Desert Valley Dolomite Cd - Cambrian Dunderberg Form ation Chp - Cambrian Highland Peak Formation

( [PLATE 2

Dsi Thrust Fault SI [ .40 ,

o p V\t IV n C ontact

Syncline

Paved or‘C d v \ ° P Highway

FIGURE 1.3: Generalized geologic map of the study area showing distribution of the Guilmette Formation (modifed from Reso, 1963; Tschanz and Pampeyan, 1970). T-4010 13

CHAPTER 2 PREVIOUS WORK

Stratigraphic Studies

The Guilmette is regionally equivalent to the Devil’s Gate Limestone of central

Nevada, the Arrow Canyon Limestone and Moapa Formation of southeastern Nevada, the

Sultan Limestone of southwestern Nevada, the Silverhom Dolomite of the Pioche District,

Nevada, the Lost Burro Formation of southern California, the Bluebell Dolomite of central

Utah, the Hyrum Dolomite of northern Utah, and the Jefferson Group and Three Forks

Formation of Idaho, Wyoming, and Montana (Poole, et al., 1977; Hintze, 1985).

The Guilmette Formation was first named and described by Nolan (1935) in

Guilmette Gulch in the Deep Creek Mountains of northwestern Utah. He described the

Guilmette as “... composed chiefly of dolomite but contains also some thick limestone beds and several lenticular sandstones.” He noted two dominant types of dolomite: a fine-grained gray dolomite containing vugs of white coarsely crystalline dolomite, and dark dolomite filled with tubular corals (Nolan, 1935). The basal contact of the Guilmette was placed immediately above the laminated dolomite of the Simonson Formation. A prominent bed containing Stri - ^ocephalus brachiopods was included within the base of the Guilmette. The upper contact was placed between a massive nonfossiliferous limestone and an overlying thin-bedded fossiliferous Mississippian limestone. Based on fossils collected, Nolan (1935) dated the Guilmette Formation as Middle to Upper Devonian.

Westgate and Knopf (1932) described a similar unit in the Silverhom Dolomite of the Bristol Range of Lincoln County, Nevada. This formation is composed dominantly of dolomite of two types: dark gray crystalline dolomite with bands of white coarsely T-4010 14 crystalline dolomite, and brown dolomite containing tubular corals replaced by white crystalline dolomite spar. Additionally, it contains dark gray limestone, sandy dolomite, and quartzitic sandstone in the upper part. The basal contact of the Silverhom was not seen, and the upper contact was placed at the base of the overlying West Range Limestone. The

Silverhom was dated as Middle Devonian, while the West Range was dated as Upper

Devonian (Westgate and Knopf, 1932). The Silverhom Dolomite is equivalent to the

Simonson and Guilmette formations, collectively (Kellogg, 1963), thus the term Silverhom

Dolomite is no longer used (Hurtubise, 1989). Neither Westgate and Knopf (1932) nor

Nolan (1935) subdivided the Guilmette and its equivalents into members.

The Guilmette has been recognized and studied in numerous ranges throughout the eastern Great Basin. Biller’s (1976) study of Upper Devonian strata in southwestern Utah included the Guilmette and its equivalents. Like Nolan (1935), he included tiieStringocephalus zone within the Guilmette, making it Middle to Upper Devonian in age. In Nevada, he placed the top of the Guilmette below the overlying West Range Limestone, but in Utah he placed the top below a sequence of siliciclastic shales, siltstones, and quartz arenites (Biller, 1976).

Kellogg (1963) included the Guilmette in his study of the Paleozoic stratigraphy of the Egan Range, Nye County, Nevada. He also placed the base of the Guilmette just above the Stringocephalus zone. Kellogg (1963) further divided the Guilmette into upper and lower members. The lower member consists of a basal massive limestone overlain by thinly bedded, yellow-weathering carbonates, and the upper member consists of interbedded cliff- forming limestones and less resistant dolomites. No sandstone was recognized (Kellogg,

1963).

Reso (1960, 1963) included the Guilmette Formation in his study of the Paleozoic stratigraphy of the Pahranagat Range, Lincoln County, Nevada. He, like Kellogg (1963), T-4010 15 placed the basal contact of the Guilmette immediately above the Stringocephalus zone. He also divided the Guilmette into lower and upper members. The lower member consists of gray limestone, yellow-weathering nonresistant dolomite, and interbedded limestone and dolomite ledges. The upper part of the lower member is a 150 foot massive cliff of limestone that Reso (1960, 1963) interpreted as biostromal in origin. The upper member contains interbedded limestone, dolomite, and sandstone. Reso (1960, 1963) placed the top of the

Guilmette between the last occurring sandstone and the overlying West Range Limestone.

Tschanz and Pampeyan (1970) included descriptions of the Guilmette Formation in their report on the geology of Lincoln County, Nevada. They placed the basal contact of the

Guilmette at the base of"... a thin-bedded, nonresistant, grayish or dusky yellow, laminated, silty dolomite” (Tschanz and Pampeyan, 1970). This places the contact 40 to 90 feet above the Stringocephalus zone. They did not divide the Guilmette into members, but recognized three interfingering facies: a sandy limestone facies, a dolomite facies, and a limestone facies (Tschanz and Pampeyan, 1970). The sandy dolomite facies is present in the

Pahranagat Range, and is equivalent to Reso’s (1960, 1963) upper member.

Hurtubise (1989) included a detailed description of the Guilmette in his study of the

Devonian system of the Seaman Range of Nye and Lincoln Counties, Nevada. He divided the Guilmette into upper and lower members, and further subdivided the lower member into the yellow slope-forming (YSF) interval and the upper interval. The basal contact of the

Guilmette was placed at the base of the distinctive YSF interval. This is above the

Stringocephalus-bearing unit, which Hurtubise (1989) named the Fox Mountain Member of the underlying Simonson Dolomite. The YSF interval “... is characterized by recessive silty to argillaceous, yellow-weathering, thin- and medium-bedded dolomite interbedded with dark gray, often well-laminated stromatolitic limestones in the basal portion” (Hurtubise, T-4010 16

1989). The upper interval of the lower member constitutes a change in weathering profile, becoming more resistant and cliff-forming than the lower YSF interval. The upper interval also contains thicker and more fossiliferous limestones than the lower YSF interval

(Hurtubise, 1989).

Hurtubise (1989) described three intervals in the upper member of the Guilmette: the lower cliff-forming interval, the middle recessive interval, and the upper cliff-forming interval. He described sandstones and sandy dolomites in the middle and upper intervals.

It is unclear, however, where he placed the boundary between the upper and lower members of the Guilmette. He placed the upper Guilmette contact at the base of the overlying West

Range Limestone, characterized by a “ . . . recessive profile of thin- and medium-bedded limestones ...” that are “ ... typically very fossiliferous, containing abundant brachiopods and less abundant gastropods, cephalopods, crinoids, and ostracods” (Hurtubise, 1989).

Ackman (1991) specifically studied the Guilmette Formation in the Schell Creek

Range and the Worthington Mountains of Lincoln County, Nevada. Like Hurtubise (1989), he placed the basal contact of the Guilmette at the base of the YSF interval, and recognized the Stringocephalus zone as the Fox Mountain Member of the Simonson Dolomite. He also placed the top of the Guilmette at the contact with the overlying West Range Limestone

(Ackman, 1991). Additionally, he divided the Guilmette into Lower, Middle, and Upper

Members. His Lower Member, like that of Hurtubise (1989) is subdivided into two intervals: the lower YSF interval and the upper ledge-forming interval. The Middle Member is only present in the Worthington Mountains, and consists of a distinctive sheer cliff bounded above and below by erosional surfaces. The Upper Member consists of the interval between the top of the Middle Member and the base of the West Range Limestone. The lithology of this member is highly variable throughout the study area (Ackman, 1991). T-4010 17

For the purposes of this study, the definitions of the basal and upper contacts of the

Guilmette and the division of the Guilmette into Lower, Middle, and Upper Members proposed by Ackman (1991) were adopted. These three members are easily defined in the

Pahranagat Range based on both lithology and depositional processes (Figure 2.1). The brachiopod Stringocephalus is an index fossil for the Givetian (Boucot, et al., 1966).

Because there is an absence of Stringocephalus in the Guilmette as defined in this study, it is thought to be completely Frasnian in age, while the underlying Fox Mountain Member of the Simonson is Givetian in age. The Middle Member occurs within the Palmatolepis punctata conodont zone (Sandberg, personal communication, 1991) which is in the Lower

Frasnian, approximately 373.5 to 373 million years before present (Ziegler and Sandberg,

1990).

Paleoecological Studies

Paleoecological and paleoenvironmental studies of the Guilmette Formation in the

Great Basin are abundant. The most comprehensive work was done by Hoggan (1975), who studied the formation in numerous ranges in eastern Nevada and western Utah. He included the Stringocephalus zone within the Guilmette, and recognized massive biostromes of bulbous stromatoporoids in the Egan Range of northeastern Nevada. Williams (1984) recognized carbonate mounds in the Guilmette in the Pequop Mountains of Elko County,

Nevada, and Smith (1984) recognized carbonate biostromal banks in the Guilmette near

White Horse Pass, Elko County, Nevada. Dunn (1979) described a prominent reef at Mount

Irish in the Timpahute Range of Lincoln County, Nevada, which is approximately 20 miles northeast of this study area. These studies are integrated into the interpretation phase of this study. T-4010 18

FIGURE 2.1: Outcrop of the Guilmette in the Pahranagat Range showing the three members described in this study: the Lower Member (A), the Middle Member (B), and the lower portion of the Upper Member (C). The base of the Guilmette occurs at the base of the first occurring yellow-weathering dolomite and is marked with a red line. Both the lower and upper contacts of the Middle Member (arrows) are erosional and represent unconformities. The Section 1 traverse is shown by a dashed line, and the area of Figure 4.9, which is a close-up, is outlined by a black box. View is to the southeast. (Photo courtesy of A.K. Chamberlain.) T-4010 19

CHAPTER 3 LOWER MEMBER OF THE GUILMETTE FORMATION

Introduction

The following descriptions and interpretations are largely based on the two strati- graphic sections measured as a basis for this report. The Lower Member of the Guilmette

Formation is easily recognized in outcrop by the presence of the characteristic yellow- weathering beds. As adopted herein, the basal contact of this member coincides with the basal contact of the Guilmette Formation, and is erosional into the underlying Fox Mountain

Member of the Simonson Dolomite (Figure 3.1). The top of this memberis placed at the base of the cliff that comprises the Middle Member. In this study the Lower Member is not further subdivided into lower and upper intervals as per Hurtubise (1989) and Ackman (1991), because that boundary could not be identified with confidence in this study area. This member is 336 feet thick at Section 1 and 411 feet thick at Section 2; however, the basal contact at Section 2 is faulted, so 411 feet is not the true thickness.

Lithology

The Lower Member consists of interbedded limestones and dolomites. Overall, the interval becomes more limy upwards. The dolomite:limestone ratio (gross thickness) near the base of the member is estimated at 9:1, becoming approximately 2:1 near the top.

Dolomite: The dolomites of the lower part of the Lower Member are gray on fresh surfaces and weather to a distinctive yellow color. These dolomites are laminated, and locally contain mudcracks, fenestrae, and tepee structures. The beds are generally 6-8 inches T-4010 20

FIGURE 3.1: Basal contact of the Guilmette Formation (arrows). The yellow-weathering dolomite is characteristic of the lower part of the Lower Member of the Guilmette. The underlying dark gray lime­ stone is the top of the Fox Mountain Member of the Simonson Dolomite. This photo was taken at Section 1 approximately 149 feet above the base. T-4010 21

thick and form recessive slopes. A second type of dolomite caps the upward-shallowing cycles in the middle and upper parts of the Lower Member, and does not weather yellow

(Figure 3.2). A third type of dolomite, referred to in this study as zebra rock, occurs

sporadically throughout the Lower Member. It is light- and dark-banded, coarsely crystal­ line dolomite that may be the result of early near-surface diagenesis. It is believed that this is similar to the banded, crystalline dolomite recognized by Nolan (1935) and Westgate and

Knopf (1932).

Limestone: The limestones of this member are gray to dark gray wackestones to packstones that form ledges. The limestones are either laminated or bioturbated and fossiliferous, with the fossiliferous units becoming more common upward. Fossils that occur in these limestones include hemispherical and bulbous stromatoporoids, the stick-like stromatoporoid Amphipora, corals, and large burrows. It is believed that the tubular corals described by Nolan (1935) and Westgate and Knopf (1932) are actually misidentified

Amphipora. Two distinctive packstone beds consisting almost entirely of oncoids (Figure

3.3) occur approximately 160 feet below the top of the Lower Member. Similar beds are also present at approximately the same level in both ranges of Ackman’s (1991) study area.

Petrography

A total of 15 thin sections from the Lower Member were examined and classified using Wilson’s (1975) Standard Microfacies (SMF). The locations of these thin sections are plotted on Plates 1 and 2. Five standard microfacies are recognized in this unit: SMF 9, SMF

16, SMF 19, SMF 20, and SMF 23. Microfacies 9 occurs in Wilson’s (1975) Facies Belt 6, which is an open marine shelf environment. The other four microfacies occur in Wilson’s T-4010 22

FIGURE 3.2: Dolomite of the medial part of the Lower Member of the Guilmette, show­ ing characteristic light color, laminations, and tepee structures (arrows). These dolomites do not weather yellow, contrasted with those in the lower part. They are generally not well exposed and form the tops of cycles. This photo is from Section 1, approximately 240 feet above the base. T-4010 23

FIGURE 3.3: The dark colored allochems in this lime packstone bed are oncoids (A). These oncoids become larger upwards, which is characteristic of such deposits. This unit sharply overlies an apparently burrowed lime wackestone (B). This photo was taken in Section 1, approximately 317 feet above the base. T-4010 24

(1975) Facies Belts 7 and 8, which are shallow platform environments.

SMF 9: characterized as bioclastic wackestone. In the Lower Member it contains fragments of brachiopods, ostracodes, gastropods, calcispheres, and conodonts that are jumbled through burrowing (Figure 3.4). An example of SMF 9 can be found in Section 2 at approximately 320 feet. As interpreted by Wilson (1975) this microfacies formed in shallow neritic water. SMF lft characterized as pelsparite/peloidal grainstone. In the Lower Member of the Guilmette the peloids may be dolomitized and are always cemented by sparry calcite.

Fossil fragments such as brachiopods, ostracodes, calcispheres, and stromatoporoids are also common (Figure 3.5). Examples of SMF 16 occur in Section 1 at 185 feet, 209 feet, and

425 feet, and in Section 2 at 360 feet. This microfacies is interpreted to represent deposition in a restricted marine shoal environment (Wilson, 1975).

SMF 19: described as laminated to bioturbatedpelletedlime mudstone to wackestone, which may grade into pelsparite with a fenestral fabric. In the Guilmette, this microfacies commonly includes ostracodes, calcispheres, and the fine-grained mud is usually dolo­ mitized (Figure 3.6). Examples of SMF 19 are found in Section 1 at 225 feet and 484 feet, and in Section 2 at 400 feet, 420 feet, and 575 feet. This microfacies is interpreted to represent deposition in restricted bays and ponds (Wilson, 1975).

SMF 20: characterized as algal stromatolite mudstone. In the Lower Member this microfacies has been dolomitized, with stromatolite layers representedby coarser crystalline dolomite (Figure 3.7). The only example of SMF 20 found in the Lower Member occurs in

Section 2 at 165 feet. This microfacies is interpreted to form in the intertidal zone (Wilson,

1975).

SMF 23: characterized as unlaminated homogeneous unfossiliferous pure micrite. T-4010 25

1.0 mm

FIGURE 3.4: Bioclastic wackestone to packstone of SMF 9. The clasts in this sample are predominantly ostracodes (A). The finer grained matrix has been almost completely dolomitized. This microfacies is interpreted to represent deposition in a subtidal shelf environment. This sample (DD 64) is from Section 2, approximately 320 feet above the base. T-4010 26

0.25 mm |

FIGURE 3.5: Peloidal grainstone characteristic of SMF 16. The dark allochems (A) are peloids and the fossil grains (B) are calcispheres. This unit has been cemented by blocky calcite spar, indicating cementation in a freshwater phreatic environment. This microfacies is interpreted to represent deposition in a restricted marine shoal environment. This sample (U2-5) is from Section 1, approximately 209 feet above the base. T-4010 27

0 .8 3 mm

FIGURE 3.6: Laminated pelleted lime mudstone of SMF 19. This sample has been partially dolomitized, as indicated by the light-colored euhedral to subhedral rhombs (A). Scattered ostracod shells are also present (B). This microfacies is interpreted to represent deposition in restricted bays and ponds. This sample (DD 80) is from Section 2, approxi­ mately 400 feet from the base. T-4010 28

I 1.0 mm

FIGURE 3.7: Stromatolite boundstone/mudstone of SMF 20. The stromatolites are indicated by the layers of slightly coarser crystalline dolomite (A). The mud has been replaced by finely crystalline dolomite. This microfacies is interpreted to have formed in the intertidal zone. This sample (DD 33) is from Section 2, approximately 165 feet above the base. T-4010 29

In the Guilmette, this microfacies is commonly dolomitized, although in at least one sample the micrite is partly altered to microspar. It may also contain windblown quartz silt (Figure

3.8). Examples of SMF 23 are found in Section 2 at 200 feet, 280 feet, and 560 feet. This microfacies is interpreted to represent deposition in highly saline tidal ponds (Wilson, 1975).

Diagenesis of the Lower Member of the Guilmette appears to have occurred soon after deposition. Due to their very finely crystalline nature, mudstones are thought to have been dolomitized penecontemporaneously. The Guilmette was deposited at a latitude of approximately 15° S, which is within the Devonian arid climate belt (Witzke and Heckel,

1988). However, little evidence of evaporite minerals has been found in the Guilmette in the study area, suggesting that it was not deposited in a hypersaline sabkha environment similar to those in the present day Persian Gulf. Rather, it may have formed in an environment similar to the Coorong region of Australia, where early dolomitization is well documented

(Dunham and Olson, 1978,1980; Muir, et al., 1980). This region is so arid that evaporation exceeds rainfall, yet is not sufficiently arid to precipitate evaporites (Muir, et al., 1980).

Ostracodes found in these mudstones are commonly filled by blocky calcite that coarsens toward the center of the paired valves. Fenestral structures appear to initially have been filled with coarsely crystalline dolomite, with remaining pore space later filled with calcite (Figure

3.9). The limestones of the Lower Member all appear to be cemented by blocky calcite spar, which implies cementation in a freshwater phreatic environment (Longman, 1982).

The juxtaposition of penecontemporaneous dolomite and freshwater calcite cement suggests a mixing-zone diagenetic environment, in which seawater mixes with freshwater to form brackish water that is saturated with magnesium and undersaturated with calcium so that dolomite forms (Badiozamani, 1973; Folk and Land, 1975). Dunham and Olson T-4010 30

0.5 mm |

FIGURE 3.8: Dolomitized micrite of SMF 23. The light-colored grains (A) are wind­ blown quartz silt. This microfacies is interpreted to represent deposition in a restricted tidal pond. This sample (DD 40) is from Section 2, approximately 200 feet above the base. T-4010 31

0.25 mm |

FIGURE 3.9: Rounded pore in a dolomitized mudstone that may be a fenestral structure. The void was intitially partially filled by large zoned dolomite rhombs (A), and then later filled by calcite (B), which has locally been dissolved, resulting in porosity (C). The zoned dolomite indicates that it formed in a mixing-zone environment. This sample (DD 56) is from Section 2, approximately 280 feet above the base. T-4010 32

(1978, 1980) showed that a mixing zone model best explains Ordovician and Silurian dolomites in central Nevada, using geochemical and paleogeographical evidence. The interbedded limestones and dolomites are the result of the fresh water table shifting up and down with changes in relative sea level (Dunham and Olson, 1978).

Environment of Deposition

The Lower Member of the Guilmette was deposited as a series of shallowing-up cycles in a shallow subtidal (below mean low tide) to supratidal (above mean high tide) environment on a broad, stable platform. In general the limestones were deposited in subtidal to intertidal zones, while the dolomites were deposited in the intertidal to supratidal zones. This interpretation is consistent with petrographic observations.

An idealized cycle in the Lower Member (Figure 3.10) has an erosional base (scour surface) filled by a thin lag. This is overlain by a fossiliferous lime packstone to grainstone

(SMF 16) unit that is usually 3 to 5 feet thick. This in turn is overlain by a mottled, burrowed, laminated lime mudstone to wackestone 5 to 7 feet thick (SMF 19), which is then overlain by a laminated dolomite mudstone (SMFs 19 and 23). This dolomite unit is generally 7 to

15 feet thick, and commonly contains such evidence of subaerial exposure as mudcracks, tepee structures, and fenestrae. Shinn (1983) commented on color variations in tidal flat deposits, with dark gray sediments representing reducing conditions in the subtidal to intertidal zone, and light gray to tan sediments representing oxidizing conditions in the intertidal to supratidal zone. The cycles of the Lower Member exhibit this variation in coloration.

The upward-increasing proportion of limestone in the Lower Member cycles indicates an overall deepening upward. Cycle thicknesses range from 3 to 41 feet, with an T-4010 33

. Scour surface filled by intraclasts- base of cycle (SMF 24) 1299^916850144 Laminated dolomite mudstone; sedimentary structures include mudcracks, fenestrae, and tepee structures (SMFs 19, 20, 23)

Burrowed, mottled lime mudstone (SMF 19)

Fossiliferous lime packstone to grainstone; fossils include stromatoporoids, corals, oncoids, brachiopods, and gastropods (SMF 16)

Scour surface filled by intraclasts- base of cycle (SMF 24)

FIGURE 3.10: Idealized cycle of the Lower Member of the Guilmette Formation. Each cycle shallows upward and may be the result of an autocyclic mechanism. The overall trend of the Lower Member is deepening-up, which may be the result of an allocyclic mechanism. T-4010 34

average cycle thickness of 15 feet. There is a minimum of 22 cycles in the Lower Member.

The bases of these cycles, interpreted as flooding surfaces, are plotted on both Section 1

(Plate 1) and Section 2 (Plate 2). The Lower Member is assumed to occupy the time between the base of the Frasnian and the base of the Palmatolepis punctata zone, thus representing approximately 1.25 million years of deposition (Ziegler and Sandberg, 1990). With a total of 22 cycles, calculated average cycle duration is approximately 56,000 years. This time estimate includes erosion and/or nondeposition associated with each cycle boundary.

Without better biostratigraphic control, however, this figure is speculative at best.

Shallowing-up cycles are common depositional features in ancient carbonate rocks

(Wilson, 1975; Laporte, 1975; Read, 1975; Shinn, 1983; James, 1984; Tucker, 1985). James

(1984) noted that “ ... carbonate accumulations repeatedly build up to sea level and above, resulting in a characteristic sequence of deposits in which each unit is deposited in progressively shallower water.” James proposed an ideal sequence within cycles, which is very similar to cycles in the Lower Guilmette. Laporte (1975) recognized three idealized facies in the Early Devonian of New York State. He described facies interpreted as supratidal, intertidal, and subtidal that are very similar to the supratidal, intertidal, and subtidal facies interpreted for the Lower Guilmette. Tucker (1985) described modem examples of shallowing-up ward cycles and further divided their environments into active tidal flat and passive tidal flat. In general, active tidal flats have more tidal channels and therefore more coarse-grained facies. The passive tidal flat sequence is dominated by pond and algal marsh deposits (Tucker, 1985). Tucker (1985) also suggested that mixing zone dolomitization is a common feature of passive tidal flats. It appears that cycles in the Lower

Member formed on a passive tidal flat. The cycles may be the result of autocyclic tidal flat progradation (Ginsburg, 1971; Tucker, 1985), while the overall deepening upward nature of T-4010 35 the member may be caused by an allocyclic mechanism such as eustasy (Grotzinger, 1986).

Goodwin et al. (1986) proposed the adoption of the punctuated aggradational cycle

(PAC) as the “fundamental unit of stratigraphic description and interpretation.” As defined, a PAC is a time-stratigraphic unit characterized as a thin (1 to 5 meters) shallowing-up cycle bounded by surfaces of nondeposition produced by a rapid rise in base level (Goodwin, et al., 1986). These cycles represent approximately a few tens of thousands of years of time and may be related to Milankovitch astronomical cycles (Goodwin, et al., 1986). According to the above definition, the cycles of the Lower Member of the Guilmette are PACs. Brett and Baird (1986) recognized PACs in the Middle Devonian of New York State, and suggest that the mechanism is “... largely facies controlled, being largely restricted to shallow shelf settings with relatively high rates of sediment accumulation.” While the approximate cycle duration of the Lower Member of 56,000 years is within the 20,000 to 100,000 year

Milankovitch band, this is not conclusive evidence for a purely allocyclic mechanism of cyclic deposition (Grotzinger, 1986).

Parasequences as defined by Van Wagoner et al. (1988) are successions of geneti­ cally-related beds bounded by marine-flooding surfaces. They are the smallest recognizable cyclic depositional sequence. The cycles of the Lower Member are generally bounded by marine-flooding surfaces and are interpreted as parasequences. T-4010 36

CHAPTER 4 MIDDLE MEMBER OF THE GUILMETTE FORMATION

Introduction

The Middle Member of the Guilmette Formation forms a distinctive cliff face that separates the less resistant Lower and Upper Members. This unit was measured as 175 feet thick at Section 1 and as 130 feet thick at Section 2. The base of this member is erosional, with up to 3 feet of relief, and locally is underlain by a lag (Figure 4.1). The top of the member is sharp, smooth, and faintly undulatory, and interpreted to represent an unconformity.

Lithology

The Middle Member has been described at Mount Irish as a breccia by Dunn (1979) and by Ackman (1991), who discussed its distribution in several ranges in southeast Nevada.

It may also be described wherever seen as a floatstone in the terminology of Embry and

Klovan (1971). It consists predominantly of limestone and dolomite clasts “floating” in a lime wackestone matrix. The dolomite clasts are generally yellow to tannish gray and laminated. The limestone clasts are predominantly whole and fragmented bulbous stromatoporoids, but also include fragments of heterolithic, laminated to bioturbated and fossiliferous (corals, stromatoporoids) limestones. In the study area the overall appearance of the Middle Member is that of dark colored clasts in a light gray matrix (Figure 4.2); however, immediately south of Section 1 (Figures 2.1, 4.9) almost the entire unit appears bedded. At Section 1, clast sizes range from as large as 4 by 60 feet to gravel sized, with clast size of about 3 to 4 inches in longest dimension most common. Clasts that are hundreds of T-4010

V ;4i

■ *

FIGURE 4.1: Basal contact of the Middle Member of the Guilmette. (arrows). This contact is erosional, and is locally underlain by a lag that is here represented by the light gray clasts. This lag was deposited separate from and previous to deposition of the Middle Member. This photo was taken approximately 100 yards south of the Section 1 traverse. T-4010 38

FIGURE 4.2: Typical clast assemblage of the Middle Member. The large, dark-colored clasts (A) are bulbous stromatoporoids, which are commonly partially silicified. The over­ all appearance of this Member is that of dark-colored clasts in a light-colored matrix. This photo was taken approximately 200 feet south of the Section 1 traverse approximately 10 feet above the base of the unit. T-4010 39 feet long have been observed in the northern Pahranagat Range, outside of the immediate areas of Sections 1 and 2 of this study (Yarmanto,inprep.). Many of the obvious larger clasts are composed of yellow-weathering dolomite similar to that found in the Lower Member

(Figure 4.3). These clasts are easily visible due to the contrast in color compared to the surrounding matrix. Large clasts are generally oriented subparallel to country-rock bedding, although several near-vertical clasts have been observed in both the Worthington Mountains

(Ackman, 1991) and the Pahranagat Range.

The upper 10 to 20 foot interval of this Member at Section 1 changes generally from matrix-supported to grain-supported, with > 50% of the grains less than 1 inch in longest dimension. Where the top of the interval has not been removed by erosion, it grades upward to a mud-sized cap. This finer-grained graded interval, which is characteristic of this

Member, has also been observed by Ackman (1991) in the Worthington Mountains, by

Chamberlain (in prep.) throughout the Timpahute Range and by Dunn (1979) atMountlrish.

Locally within the Pahranagat Range this bed has been variably removed by erosion.

From a distance, the entire unit appears bedded in both Section 1 and Section 2. At

Section 2 bedding planes can be traced across the entire outcrop, an approximate distance of one half mile. Locally bulbous stromatoporoids are concentrated in lens-shaped beds

(Figure 4.4). An oriented hand sample from approximately 250 feet south of the Section 1 traverse collected from about 100 feet above the base of the Middle Member is a stromatoporoid boundstone. Additionally there is what appears to be a stromatoporoid biostrome with stromatoporoids in growth position (Figure 4.5) that is located approx­ imately 100 yards southeast of the Section 1 traverse 25 feet from the top of the Member.

This biostrome occurs within apparently cyclic limestones separated by bedding planes parallel to regional strike and dip. The distribution of lithofacies within the Middle Member T-4010 40

FIGURE 4.3: An example of very large clasts within the Middle Member. These clasts, which are outlined in black, are composed of thinly-bedded, yellow-weathering dolomite that is similar to the dolomite of the underlying Lower Member. The large clast at the bot­ tom of the picture is approximately 5 feet high. This photo was taken approximately 400 feet north of the Section 1 traverse, approximately 10 feet above the base of the Member. T-4010 41

FIGURE 4.4: Lens-shaped bed of bulbous stromatoporoids. Beds such as this are found throughout the Middle Member, and are easily recognized because the stromatoporoids are partially silicified. This photo was taken on the Section 2 traverse, approximately 20 feet above the base of the unit. T-4010 42

FIGURE 4.5: Biostrome of stromatoporoids in growth position (arrows). The growth morphology of these fossils indicates that they were trying to keep pace with rising sea level. This biostrome occurs within a sequence of limestone beds containing Amphipora and corals. This photo was taken approximately 100 yards south of the Section 1 traverse 25 feet below the top of the Member. T-4010 43 has not been mapped. It is possible that some of the samples in this study came from matrix, some from small clasts, and some from large clasts whose margins are not obvious.

Petrography

A total of eight thin sections from the Middle Member were examined and classified using Wilson’s (1975) Standard Microfacies (SMF). The approximate footages of these thin sections were projected onto the traverse of Section 1 and are shown on Plate 1. Seven of these thin sections were taken from samples interpreted at the time of sampling to represent matrix. Five standard microfacies are recognized in this unit: SMF 4, SMF 5, SMF 16, SMF

20, and SMF 24. Standard Microfacies 4 and 5 are found in Wilson’s (1975) Facies Belts

3 and 4, which are interpreted as relatively deep water slope environments. Standard

Microfacies 16, 20, and 24 are found in Wilson’s (1975) Facies Belts 7 and 8, which are interpreted as shallow platform environments.

SMF 4: characterized as microbreccia or bioclastic-lithoclastic packstone. In the

Middle Member of the Guilmette, the grains are polymictic and include previously- cemented lithoclasts and bioclasts containing fossil grains such as Amphipora, brachio- pods, and bivalves (Figure 4.6). Samples JE 91-4 from Section 1 and DD 120 from Section

2 are interpreted to represent this microfacies, which is commonly cemented by coarse equant calcite and is interpreted to represent deposition in a slope environment above storm wave base (Wilson, 1975).

SMF 5: characterized as bioclastic grainstone-packstone or floatstone. In the

Guilmette, this microfacies differs from SMF 4 in that the grains are “floating” in the matrix, and at least 50% of the grains are bioclastic in origin. In this sample (JE 91-5), which was collected from matrix, the bioclasts are predominantly Amphipora. The matrix has been T-4010 44

1.0 mm

FIGURE 4.6: Bioclastic-lithoclastic packstone of SMF 4. The large grain at the top of the photo (A) is a fragment of Amphipora, while under it (B) is a lithoclast of previously- cemented peloids. This photo also exhibits the relatively high interparticle porosity of the Middle Member. This sample (JE 91-4) was taken from approximately 657 feet above the base of Section 1, near the top of the Middle Member. T-4010 45

completely dolomitized, while the allochems have remained calcitic (Figure 4.7). This microfacies is found in the upper two feet of the Middle Member at Section 1, and is interpreted to represent deposition in a slope environment above storm wave base (Wilson,

1975). Both SMFs 4 and 5 are thought to be the most characteristic of the matrix of the

Middle Member.

SMF 16: described as pelsparite/peloid grainstone. In this sample (JE 90-16)

allochems are dominantly peloids, with minor amounts of ostracodes, brachiopods, and calcispheres. Grains are cemented by sparry calcite. This sample was collected from a yellow-weathering dolomite clast approximately 20 feet above the base of the Middle

Member at Section 1. This clast is interpreted to have come from the underlying Lower

Member, an interpretation which is supported by the thin section analysis. This microfacies is identical to SMF 16 in the Lower Member, and is interpreted to have originally been deposited in a restricted marine shoal environment (Wilson, 1975).

SMF 20: characterized as algal stromatolite mudstone. In the Middle Member of the Guilmette, this microfacies occurs in two samples with enigmatic laminations that were collected from Section 1. One was collected at the base of the Middle Member 100 yards south of the traverse (JE 91-1 A); the other (JE 91-2) was collected approximately 250 feet north of the traverse at a prominent surface approximately 80 feet from the base of the

Member. The samples are wackestones to packstones, with allochems of previously cemented lithoclasts. These allochems appear to be bound or trapped by vague laminations

(Figure 4.8), which are probably algal in origin. This microfacies is interpreted to occur in the upper intertidal zone (Wilson, 1975).

SMF 24: described as coarse lithoclastic-bioclastic rudstone or floatstone. In the

Middle Member, this sample (JE 91-IB) was collected from a lag 2 to 3 inches thick just T-4010 46

1.0 mm

FIGURE 4.7: Bioclastic floatstone of SMF 5. This sample (JE 91-5) was taken from approximately 660 feet above the base of Section 1, which is at the top of the Middle Member. The bioclasts are predominantlyAmphipora fragments (A). The matrix has been completely replaced by euhedral to subhedral dolomite (B). The majority of the clasts have remained calcitic, indicating that they were well lithified prior to dolomitiza- tion.

ARTHUR LARES LIBRARY COLORADO SCHOOL OF MINU GOLDEN, CO 80401 T-4010 47

1.0 mm

FIGURE 4.8: Stromatolitic mudstone-boundstone of SMF 20. This sample (JE 91-2) was taken approximately 250 feet north of the Section 1 traverse from a prominent sur­ face approximately 80 above the base of the Middle Member. The stromatolites appear to be binding the finer grained sediment, and trapping the scattered windblown quartz grains (A). This microfacies is interpreted to form in intertidal environments. T-4010 48 below the base of the Member. Lithoclasts are predominantly micrite or previously- cemented peloid grainstones that are coated with algal laminations. The matrix is fine­ grained limestone. This microfacies is interpreted to represent deposition as a lag (Wilson,

1975) above an erosional surface .

In all of the samples except SMF 5 cement is dominantly sparry equant calcite, which may form either in a freshwater phreatic environment or a deep burial environment

(Longman, 1982). Cloudy dolomite rhombs are sparsely scattered throughout all of the samples examined from this Member. In the sample of SMF 5 (JE 91-5) the matrix has been completely replaced by finely crystalline dolomite, while the previously-cemented clasts have remained calcitic. Because the Middle Member is interpreted to have formed in a relatively deeper marine environment, it could be expected to exhibit submarine phreatic cement; however, no such cements have been observed. Finely disseminated pyrite observed in nearly all of the samples indicates a reducing diagenetic environment for at least part of its history.

A relative sea level lowstand is interpreted to have followed the deposition of the

Middle Member, thus causing it to be flushed with meteoric water. This action may have resulted in the dissolution of the higher Mg-calcite marine cement and reprecipitation of freshwater calcite cement. The mobilized Mg may have been sufficient to locally precipitate dolomite as rhombs sparsely scattered throughout the Middle Member. Because the matrix became completely dolomitized at the top of the Member, it was probably subjected to prolonged exposure in a mixing zone environment. The majority of fractures observed in thin section were filled by coarse equant calcite. These fractures probably resulted from tectonism, and were later filled during post-uplift meteoric flushing. T-4010 49

Environment of Deposition

The exact origin of the Middle Member is enigmatic. There are many unanswered questions about the origin of this Member, and a complete interpretation is beyond the scope of this study. The problem is currently being studied by Alan Chamberlain, Yarmanto, and

John Warme at Colorado School of Mines.

Hoggan (1975) described breccias within the Guilmette in the Leppy Range, Snake

Range, and Egan Range, and attributed them to solution-collapse phenomena, and Larsen, et al. (1988) mentioned “polymict to monomict collapse breccias” from western Utah. It is not known how or if these breccias correlate to the Middle Member, and no evidence of karst or collapse processes were observed in the study area.

The predominance of well-preserved stromatoporoid clasts suggests proximity of biostromes to the study area; however, the only obvious reef recognized, to my knowledge, in the Guilmette thus far occurs at Mount Irish, in the Timpahute Range, immediately above the Middle Member (Dunn, 1979)

In both measured sections of the study area this Member appears bedded (Figure 4.9).

Additionally, there appear to be at least four bedsets, which are separated by surfaces that are recessive into the otherwise shear cliff face (Figure 4.9). The apparently bedded nature of this Member in this location is anomalous when compared to the Middle Member in other locations in the Pahranagat Range and elsewhere (Ackman, 1991; Chamberlain, in prep.;

Yarmanto, in prep.), and is difficult to trace laterally.

Warme et al. (1991) suggested that this Member was deposited as the result of a single, catastrophic event, possibly owing to energy from the impact of an extraterrestrial object into the earth. This interpretation assumes that the entire Middle Member is one bed that generally fines upward. Large, partially attached clasts that Ackman (1991) termed T-4010 50

FIGURE 4.9: Oblique aerial view of the Middle Member of the Guilmette, showing the apparently bedded nature of the unit. This photo was taken from a helicopter approxi­ mately 75 yards south of the Section 1 traverse. The arrows indicate two of the recessive surfaces that separate apparent bedsets. Approximate distance between the two surfaces is 35 feet. (Photo courtesy of A.K. Chamberlain.) T-4010 51

“peel structures” were noted in the Worthington Mountains and in the Pahranagat Range north of the study area. The presence of such clasts, as well as other very large clasts, supports the proposed catastrophic origin of the Middle Member. Warme (personal communication, 1991) suggested that the four bedsets recognized in the Pahranagat Range in the vicinity of Section 1 (Figure 4.9) are actually large bedded clasts stacked one on top of another, perhaps even the same bedset broken and stacked in an imbricated fashion. T-4010 52

CHAPTER 5 UPPER MEMBER OF THE GUILMETTE FORMATION

Introduction

The basal contact of the Upper Member is placed at the unconformity that marks the top of the Middle Member. The upper contact of this member coincides with the upper contact of the Guilmette Formation, and is placed at the base of the overlying Pilot Shale.

This member was measured as 1738 feet thick at Section 1 and was not measured at Section

2. This member is believed to correlate, at least in part, with the Upper Member described by Reso (1960, 1963), Larsen, et al. (1988), Hurtubise (1989), and Ackman (1991).

Elsewhere in the Pahranagat Range, Reso (1960) measured the Upper Member as 1634 feet at Lower Downdrop Mountain (Sections 19 and 20, T 6 S, R 59 E), 1572 feet at Downdrop

Mountain Repeat (Sections 21,28, and 29, T 6 S, R 59 E), and 1639 feet at Devonian Ridge

(Sections 15 and 16, T 7 S, R 59 E).

Lithology

The Upper Member is composed of interbedded limestones, dolomites, and terri­ genous sandstones. It is the high percentage of quartzose sandstone that makes this

Pahranagat section of the Guilmette unique compared with most other published sections.

Limestone: The limestones of this member were subdivided into three general types: laminated mudstone, bioturbated wackestone, and fossiliferous packstone. The laminated mudstones are gray, generally unfossiliferous, and form recessive slopes. The bioturbated wackestones are dark gray and mottled, with beds 6 inches to 5 feet thick. T-4010 53

Burrows up to 6 inches long are the dominant feature of this facies (Figure 5.1). Scattered

gastropods and brachiopods are also present. The fossiliferous packstones are dark gray and form thick (5 to 10 feet) ledges. Fossils commonly present in this facies include

stromatoporoids, stromatolites, corals, gastropods, brachiopods, and megalodont bivalves

(Figure 5.2).

Dolomite: The dolomites of the Upper Member may also be subdivided into three types: laminated mudstone to wackestone, recrystallized fossiliferous limestone, and zebra rock. The laminated mudstones to wackestones weather to light gray, contain no fossils, locally exhibit tepee structures and windblown sand, and form recessive, generally covered, slopes. The fossiliferous limestones that are recrystallized to dolomite weather to light gray and generally form massive cliffs. In some cases recrystallization has obscured the original fossils. In most cases, however, the fossils are easily recognized and most commonly are stromatoporoids, including the genus Amphipora. Less abundant fossils include gastropods, brachiopods, and corals. The zebra dolomites consist of alternating light and dark gray bands of coarsely crystalline dolomite and are similar to the zebra dolomites described in the Lower Member. Because this fabric is parallel to bedding it is interpreted to result from early near­ surface diagenesis.

Sandstone: The sandstones of the Upper Member are all classified as quartz arenites because they are composed of >95% quartz grains (Pettijohn, et al., 1987). There are six general types of sandstone. The most commonly occurring type of sandstone is massive and homogeneous with no apparent sedimentary structures. Upon close inspection these sandstones appear to have been bioturbated, and locally planar horizontal laminations are preserved. These sandstones are generally fine- to medium-grained, well rounded and well sorted. T-4010 54

FIGURE 5.1: Bioturbated lime wackestone of the Upper Member of the Guilmette. This photo was taken from approximately 670 feet above the base of Section 1. View is perpendicular to bedding plane. Burrows such as these are characteristic of this wacke­ stone facies. T-4010 55

FIGURE 5.2: Fossiliferous lime packstone of the Upper Member of the Guilmette. On- coids (A) and megalodont bivalves (B) are common fossils of this facies. This photo was taken from Section 1 at approximately 1680 feet. This unit is interpreted to have formed in a restricted lagoon environment. T-4010 56

The next most common type of sandstone is burrowed, and differs from the first type in that the individual burrows are preserved, the grain size is very fine to fine, and the individual beds fine upward. A third type of sandstone is trough cross-bedded and forms very prominent cliffs. Troughs are generally 6 to 8 inches high and 10 to 12 inches wide, with a westerly component of paleotransport direction (Figure 5.3). These sandstones commonly have scoured bases, are fine-grained, well rounded and well sorted. Sandstones of the fourth type are similar to those of the first type in being fine- to medium-grained, well rounded and well sorted, but appear to be irregularly bound by algal mats (Figure 5.4). This facies is commonly associated with zebra dolomite and locally contains oncoids and brecciated clasts.

The fifth type of sandstone forms poorly exposed, slope-forming units that are planar and/or ripple laminated, and locally bioturbated. These sandstones vary from fine- to coarse­ grained and locally contain up to 10% lithic fragments. There is only one occurrence of the sixth type of sandstone, which is a fine- to medium-grained, very well rounded, very well sorted quartz arenite. Its most notable characteristic is the large scale planar-tabular cross- bedding (Figure 5.5). This type occurs in Section 1 at 1985 to 2025 feet, is approximately

40 feet thick, has a sharp base, and is capped by a laminated limestone.

Petrography

A total of 38 thin sections from the Upper Member were examined and classifed using Wilson’s (1975) Standard Microfacies (SMF). The locations of these thin sections are plotted on Plate 1. Four existing standard carbonate microfacies (SMF 16, SMF 19, SMF

20, and SMF 23) were recognized, and one additional microfacies (SMF 25) was created to describe the sandstone units. All of these microfacies are interpreted to represent deposition T-4010 57

FIGURE 5.3: Longitudinal view of trough cross-beds commonly found in sandstones of the Upper Member. In this photo, which was taken from Section 1 at approximately 810 feet, structural dip has been rotated out, showing true paleocurrent direction (arrow) is from northeast to southwest. These cross-beds may have formed as large ripples in a tidal channel. T-4010 58

FIGURE 5.4: This type of sandstone commonly found in the Upper Member appears to be irregularly bound by algal mats (arrows). Sandstones such as this one probably deve­ loped on exposed point bars and/or levees of tidal channels. This photo was taken in Section 1 at approximately 1265 feet. T-4010 59

FIGURE 5.5: The most notable characteristic of this sandstone from the Upper Member is the planar-tabular cross-bedding. It is interpreted to represent an eolian dune that mi­ grated across the tidal flat. The only occurrence of this type of sandstone in Section 1 is located between 1985 and 2025 feet. T-4010 60 in a shallow subtidal to supratidal environment (Wilson, 1975).

SMF 16: characterized as pelsparite/peloidal grainstone. In the Upper Member the peloids are cemented by sparry calcite. In some samples an early microspar cement was followed by later coarse equant calcite. Fossils such as gastropods, calcispheres, and nodosinellid forams (Figure 5.6) are presentin this microfacies. Examples of SMF 16, which is interpreted to represent deposition in a restricted marine shoal environment (Wilson,

1975), are found in Section 1 at 730 feet, 1525 feet, 1665 feet, 1905 feet, 1945 feet, and 2285 feet.

SMF 19: described as laminated to bioturbatedpelletedlime mudstone to wackestone.

In the Upper Member, the fine-grained mud is commonly dolomitized, and isolated quartz grains are present along individual laminae (Figure 5.7). Examples of SMF 19 are found in

Section 1 at 1725 feet, 1925 feet, and 2125 feet, and one sample is found in Section 2 at 740 feet. This microfacies is interpreted to represent deposition in restricted bays and ponds

(Wilson, 1975).

SMF 20: described as algal stromatolite mudstone. Two examples of this microfacies, located in Section 1 at 925 feet and 2164 feet, were recognized in the Upper

Member. This microfacies composed of alternating micrite and stromatolitic laminations

(Figure 5.8). The stromatolites have been preferentially dolomitized. This microfacies is interpreted to represent deposition in the intertidal zone (Wilson, 1975).

SMF 23: characterized as unlaminated homogeneous unfossiliferous pure micrite.

In the Guilmette this microfacies is commonly dolomitized (Figure 5.9). All but one of the samples of this microfacies contain up to an estimated 20% quartz grains that are generally fine- to medium-grained, and moderately rounded and sorted. Examples of SMF 23 are found in Section 1 at 1085 feet, 1105 feet, 1205 feet, 1542 feet, 1585 feet, 1925 feet, and 2385 T-4010 61

FIGURE 5.6: Peloidal grainstone of SMF 16. Common allochems in this microfacies are peloids (A), calcispheres (B), and forams (C). This slide appears to show two epi­ sodes of cementation: an early microspar calcite and a later coarse equant calcite. This sample (U12-7) was taken from Section 1 at approximately 1905 feet and is interpreted to represent deposition in a restricted marine shoal environment. T-4010 62

FIGURE 5.7: Pelleted lime mudstone to wackestone of SMF 19. This sample (U12-8) from 1925 feet in Section 1 is generally unfossiliferous, and locally is up to 50% dolomi­ tized. The large grains (A) are fine-grained, windblown quartz sand. This microfacies is interpreted to represent depostion in restricted bays and ponds. T-4010 63

0.5 mm

FIGURE 5.8: This sample (U14-1) is a good example of SMF 20, described as algal stromatolite mudstone, which in this case has been dolomitized. The stromatolite lami­ nae (A) have been replaced by dolomite crystals that are larger than the surrounding finely crystalline dolomite mudstone. This microfacies is interpreted to form in the inter­ tidal zone. This sample was collected from 2164 feet at Section 1. T-4010 64

0.5 mm |

FIGURE 5.9: This sample (U12-8) was collected from 1925 feet at Section 1 and is a typical example of SMF 23. In this slide the micrite has been completey replaced by subhedral dolomite. This microfacies is interpreted to represent deposition in intermit­ tent, saline tidal ponds. T-4010 65

This microfacies is interpreted to represent deposition in highly saline tidal ponds (Wilson,

1975). The quartz was deposited as windblown sand into these intermittent ponds.

SMF 25: described as quartz arenite with dolomite cement. The monocrystalline quartz grains vary from fine- to coarse-grained, subangular to very well rounded, and poorly to well sorted. In many samples the quartz grains have polygonal boundaries, indicating a phase of quartz overgrowth cementation (Figure 5.10). The dolomite cements are coarsely crystalline with an equant mosaic texture. Examples of SMF 25 are found in Section 1 at 805 feet, 905 feet, 985 feet, 1145 feet, 1265 feet, 1540 feet, 1602 feet, 1805 feet, 2005 feet, 2145 feet, 2230 feet, 2305 feet, 2345 feet, and 2365 feet. This microfacies is inteipreted to represent quartz sand deposition in shallow subtidal to supratidal environments.

A variety of diagenetic environments are recognized in the Upper Member of the

Guilmette. The fine-grained dolomites of SFM 19, SMF 20, and SMF 23 were dolomitized in amixing-zone environment soon after deposition, as discussed in the chapter on the Lower

Member. The peloid grainstones of SMF 16 were cemented by blocky, sparry calcite in a freshwater phreatic environment (Longman, 1982). The coarsely crystalline dolomite samples indicate later secondary dolomitization, in which limestone has been completely recrystallized to dolomite. This process commonly happens during later burial (Longman,

1982). The sandstones commonly exhibit two stages of cementation: an early quartz overgrowth phase and a later dolomite phase. Many of the samples exhibit fractures that are filled with equant calcite. These represent the final phase of diagenesis, which is probably related to Tertiary tectonic activity.

Environment of Deposition

In general, the Upper Member of the Guilmette was deposited in shallow subtidal to T-4010 66

0.5 mm

FIGURE 5.10: This microfacies, SMF 25, was created to describe sandstones in the Upper Member of the Guilmette. This slide shows a quartz arenite with two stages ol cementation: an early quartz overgrowth phase (A) and a later, pore-filling dolomite phase (B). This microfacies represents a variety of peritidal environments. T-4010 67

supratidal environments. The presence of siliciclastic sediments interbedded with carbonate

sediments suggests an active tidal flat environment (Tucker, 1985). This member appears

to be somewhat cyclic, although the cycles are not as well documented as the cycles in the

Lower Member. Less field time was spent describing and analyzing the Upper Member due

to its inaccessibility. Because the cycles are so variable an idealized cycle is difficult to construct. Some common facies associations that occur in somewhat cyclical sequences are

illustrated in Figure 5.11. Cycle boundaries are picked at the contact of a subtidal facies overlying a shallower facies. The interpreted cycle bases are plotted on Section 1 (Plate 1).

The Upper Member is assumed to occupy the time between the top of the Palmatolepis punctata zone and the top of the Frasnian, thus representing approximately 3 million years of deposition (Ziegler and Sandberg, 1990). Forty-seven cycles were recognized in the

Upper Member, making the average cycle duration approximately 63,000 years, which is within the Milankovitch band of20,000 to 100,000 years. The cycles are probably the result of both autocyclic and allocyclic mechanisms.

In general, the burrowed lime mudstones and the type A and B sandstones are interpreted as deposited in subtidal environments. The burrowed lime mudstones (SMF 19) are interpreted to represent deposition in restricted lagoons, while the peloidal grainstones

(SMF 16) and sandstones are interpreted to represent deposition in more normal marine environments. The fossiliferous packstones, both limestone and dolomite, are interpreted to represent deposition in subtidal to intertidal zones. The fossil content indicates a restricted, somewhat quiet environment (Hoggan, 1975), such as lagoons behind barriers.

Facies associations suggest that type E sandstones were deposited in an intertidal environ­ ment, such as a low energy swash zone.

Type C sandstones are interpreted to represent deposition in tidal channels that cut T-4010 68

<3> Fossiliferous packstone- a grainstone P <> -Cycle boundary Laminated mudstone <3> Fossiliferous packstone- a p O grainstone “Cycle boundary

An example of this facies association occurs between 2030 and 2100 feet at Section 1.

Trough cross-bedded sandstone

Cycle boundary Zebra dolomite Algal-bound sandstone

Trough cross-bedded sandstone .Cycle boundary

An example of this facies association occurs between 1225 and 1290 feet at Section 1.

FIGURE 5.11: Common cyclical facies associations in the Upper Member of the Guilmette. These cycles are not as well-developed as those of the Lower Member, and are probably the result of both autocyclic and allocyclic mechanisms. T-4010 69 across the predominantly carbonate tidal flats. Hoggan (1975) suggested a cratonic source area to the east, and paleocurrent data from this study area support this interpretation. The composition and texture of the sands suggest a distant source or multi-cycle provenance.

While no provenance analyses were performed, Lower Paleozoic sandstones, such as the

Eureka Quartzite, on small local uplifts are a likely source. Biller (1976) recognized cross­ bedded quartz arenites similar to type C sandstones in his measured sections of the Guilmette in the Pavant Range, the Confusion Range, the North Gilson Mountains, and the South

Burbank Hills, Utah. He described these sandstones as forming lenticular ledges that grade laterally into quartzose carbonate rocks. The type D algal-bound sandstones are interpreted as intertidal to supratidal deposits and may have formed on bars and levees of tidal channels.

Laminated unfossiliferous limestones and dolomites (SMF 23) indicate intertidal to supratidal deposition. The type F sandstone may be a relict eolian dune that migrated across the tidal flats and was eventually buried by the transgressing sea. Modem analogs were described by Phleger and Ewing (1962) in Baja California and by Shinn (1973) on the Qatar

Peninsula of the Persian Gulf. The dolomite cement present in all of the sandstones sampled for thin sections supports the tidal flat depositional model (Shinn, 1973).

Klein (1977) stated that interbedded carbonates and quartz arenites are “... common to stable craton^. or the epeiric, platform or mioclinal shelf sea association” but few ancient examples of this association have been documented. Driese and Dott (1984) described similar cycles in the Middle Pennsylvanian Morgan Formation and proposed a two-phase depositional model. The first phase consists of instantaneous transgression followed by progradational carbonate deposition. The second phase consists of progradation of sand­ stones across intertidal and subtidal carbonate rocks. Eustatic changes related to glaciation are thought to be the mechanism that generated these cycles (Driese and Dott, 1984). T-4010 70

Cycles described by Larsen, et al. (1988) in the Upper Guilmette in Utah are very similar to the carbonate cycles of the Upper Guilmette in the Pahranagat Range. Larsen, et al. (1988) also recognized an overall transgressive trend of the Upper Guilmette to the contact with the overlying deep-water Pilot Shale. The Pahranagat section does not follow this trend, however. It is transgressive only above approximately 2025 feet on Section 1

(Plate 1). Below 2025 feet the unit is interpreted as regressive due to the predominance of intertidal to supratidal quartz arenites. In this case the relative sea level changes could be directly related to local tectonics, as indicated by the presence of the sandstones. T-4010 71

CHAPTER 6 SEQUENCE STRATIGRAPHIC FRAMEWORK OF THE GUILMETTE FORMATION

Introduction

Sequence stratigraphy is defined by Van Wagoner, et al. (1988) as “ ... the study of rock relationships within a chronostratigraphic framework of repetitive, genetically related strata bound by surfaces of erosion or nondeposition, or their correlative conformities.” To date, no detailed sequence stratigraphic interpretation of the Guilmette has been published.

The limited areal extent of this study precludes a regional interpretation, and the lack of accessible dip exposures near the study area limits the accuracy of the interpretation offered herein. The Guilmette strikes approximately N15°W in the Pahranagat Range. To the west of the study area, the Guilmette crops out in the JumbledHills which are within the Las Vegas

Bombing and Gunnery Range, and access there is severely restricted and not available for this study. The Pahroc Range to the east of the study area contains no Guilmette exposures.

The closest Guilmette exposures to the east occur in the South Burbank Hills and Blawn

Mountain, both in Utah. To the southeast the Guilmette crops out in the Delamar Range. It is hoped that the observations made in this study will be useful to those attempting to interpret the Guilmette regionally.

Key Definitions

The following discussion is summarized from Van Wagoner, et al. (1988) and Sarg

(1988). A sequence is a succession of genetically related strata bounded by unconformities and their correlative conformities. Sequences may by subdivided into systems tracts, which T-4010 72 are classified based on their position in the sequence. Systems tracts are composed of parasequence sets, which are in turn composed of parasequences. A parasequence is a succession of genetically related beds bounded by marine-flooding surfaces. Parasequences are the smallest recognizable cyclic depositional sequence. The cycles of the Lower Member are considered parasequences.

The transgressive-systems tract (TST) is characterized by one or more landward- stepping parasequence sets deposited during a relative rise in sea level. The top of the TST is the surface of maximum flooding. The highstand-systems tract (HST) consists of aggradational (vertically stacked) to progradational (seaward-stepping) parasequence sets deposited during maximum sea level rise and/or sea level stillstand. The HST is bounded below by the surface of maximum flooding and above by either a type 1 or type 2 sequence boundary. A type 1 sequence boundary is characterized by subaerial exposure and concurrent subaerial erosion accompanied by a basinward shift in facies, resulting in nonmarine strata being deposited over deeper marine facies. In carbonate systems, type 1 sequence boundaries form when sea level drops below the shelf edge and the entire shelf is subaerially exposed. A type 2 sequence boundary is marked by subaerial exposure that lacks subaerial erosion and basinward facies shift. In carbonate systems type 2 sequence boundaries form during small-scale falls in sea level, in which only the inner platform is exposed. The shelf margin wedge is deposited over type 2 sequence boundaries during the following rise in sea level.

Fischer Plots

A Fischer plot is a graph in which cumulative cycle thickness of peritidal carbonates is corrected for linear subsidence and plotted versus time using an average cycle period T-4010 73

(Read and Goldhammer, 1988). Read and Goldhammer (1988) used Fischer plots to define changes in relative sea level at the third order scale in Ordovician carbonates in the

Appalachians. The model assumes that long term rise in sea level (or increase in subsidence rate) will produce thick peritidal cycles, while long term fall in sea level (or decrease in subsidence rate) will produce thin cycles (Read and Goldhammer, 1988).

An example of a hypothetical Fischer plot is presented in Figure 6.1. This graph was plotted using a computer program developed by Gary Vander Stoep of the Bureau of

Economic Geology at the University of Texas. This program builds a database of average cycle duration, average subsidence per cycle, and cycle thickness for each cycle, and then plots the data on a graph. Average subsidence per cycle is calculated by dividing total thickness of the unit by number of cycles in the unit. On this graph, the vertical distance between cycle tops represents the average cycle duration, which is calculated by dividing total time of the unit by the number of cycles in the unit. The path of linear subsidence results from plotting cumulative cycle thickness versus average cycle duration. Connecting the tops of each cycle results in a third order sea level curve for the unit (Read and Goldhammer,

1988). The Fischer plot method was used for this study because it appears to be an objective means of tracking sea level changes. The weakness in this method is the assumption of linear subsidence, which appears to favor eustatic over tectonic mechanisms. This method is not accurate for the Middle Member because it is not peritidal, and may be less accurate for the

Upper Member because of the clastic influx. Average cycle thickness and duration were calculated separately for each Member.

A column showing assumed age of the Guilmette Formation is presented in Figure

6.2. The only unit that has been dated with biostratigraphy is the Middle Member, which occurs entirely within the Palmatolepispunctata conodont zone (Sandberg, 1991, personal 10 o -10 Cumulative Thickness (feet)

FIGURE 6.1: Hypothetical Fischer plot showing the components of the curve. The distance between each cycle top is equal to the average cycle period. Path of linear subsidence is calculated from average cycle duration vs. cumulative cycle thickness. Cycle thickness is calculated from average subsidence per cycle. Connecting the cycle tops generates a third order sea level curve (Read and Goldhammer, 1988). T-4010 75

<

LU Age (million years) < LL 370

oz DC LU LL Q t= Palmatolepis punctata 3 7 3 LU Middle Member

3 Palmatolepis punctata 373.5 ( 3 Lower 374 Member

375

LU > o

FIGURE 6.2: Stratigraphic column showing age of the Guilmette Formation as interpreted in this study. The Middle Member was dated using conodont data; the other units were not dated (modified from Geological Society of America, 1983; Ziegler and Sandberg, 1990). T-4010 76 communication). This zone, which is equivalent to the former Middle asymmetricus zone, is estimated to be 500,000 years in duration (Ziegler and Sandberg, 1990). For the purpose of this exercise this Member is assumed to have five cycles, each averaging 100,000 years in duration. The Lower Member is assumed to represent the span of time between the base of the Frasnian and the base of the Palmatolepis punctata zone, thus representing approxi­ mately 1.25 million years of deposition. There are 22 cycles in the Lower Member, and average cycle duration is approximately 56,000 years. The Upper Member is assumed to occupy the span of time between the top of the Palmatolepis punctata zone and the top of the Frasnian, thus representing approximately 3 million years of deposition. There are 47 cycles in the Upper Member, and average cycle duration is 63,000 years. The computer­ generated Fischer plot of the Guilmette at Section 1 in the Pahranagats is presented in Figure

6.3.

Interpretation

The relative sea level curve constructed from the Fischer plot for the Guilmette is shown in Figure 6.4. The Lower Member shows an overall transgressive, landward-stepping trend and is interpreted to represent the transgressive-systems tract. The erosional surface which marks the contact between the Lower and Middle Members represents the surface of maximum flooding and may be an unconformity. The Middle Member is interpreted as representing the highstand-systems tract, because the accommodation space necessary for this deposit to accumulate requires a high stand of relative sea level. The upper contact of the Middle Member is interpreted to represent a relative sea level fall because the unit immediately overlying it is composed of laminated dolomite interpreted as supratidal due to the presence of tepee structures and mud cracks. Additionally, the pervasive dolomitization T-4010 77

Upper Member

2 Mid. Mei

1 Lower Member

0 300 200 100 0 -100

Cumulative Thickness (feet)

FIGURE 6.3: Fischer plot of the Guilmette Formation as measured at Section 1 in the Pahranagat Range. The curve for each Member was calculated separately, based on average cycle thickness and duration for each Member. The Middle Member is assumed to have five cycles, based on average cycle duration calculated for the other two Members. T-4010 78

Relative Sea Level Rise Fall

5 - Shelf Margin Upper Member Wedge 4 -

C/) TO Q— >O ° c o

iequencej Boundary Highstand Mid. Mem Systems Tract

I Trans­ 1- Lower gressive Member Systems i T ract

+100 0 -100 (Feet)

FIGURE 6.4: Relative sea level curve for the Guilmette Formation in the Pahranagat Range, with Members and their interpretations delineated. This 3rd order curve was plotted from the high points of each cycle on the Fischer plot in the previous figure. T-4010 79 of the matrix at the top of the Middle Member may indicate prolonged exposure to a mixing zone environment. Sarg (1988) noted that mixing zone dolomitization is common below type 2 sequence boundaries; therefore the top of the Middle Member is interpreted as such a surface. This fall is indicated by a small regressive pulse on the Fischer plot.

The Fischer plot of the Upper Member shows a short relative sea level rise followed by a long term sea level fall. The upper one-third of the Upper Member shows a small transgression followed by a regression. These short term pulses may be related to local tectonics, while the long term regression may be a regional event. The Upper Member is interpreted as a shelf margin wedge deposited above the type 2 sequence boundary.

Johnson and Sandberg (1988) published a relative sea level curve for the Devonian in the western United States. According to their curve, the Lower Member of the Guilmette as described in this study is equivalent to the Lowermost asymmetricus-zone eustatic rise.

The Middle Member is equivalent to the Middle asymmetricus-zone eustatic rise, which produced biohermal buildups in the Devil’s Gate Formation. The base of the Upper Member is a continuation of the Middle asymmetricus-zone eustatic rise. Using the Johnson and

Sandberg (1988) curve, the entire Guilmette could be interpreted as one transgressive- systems tract. T-4010 80

CHAPTER 7 PETROLEUM POTENTIAL OF THE GUILMETTE FORMATION

Introduction

The Guilmette Formation is a proven hydrocarbon reservoir at both Bacon Flat and

Grant Canyon Fields in Railroad Valley, Nye County, Nevada (Veal, et al., 1988). Bacon

Flat Field has produced over 267,000 barrels of oil since its discovery in 1981 (Veal, et al.,

1988). Grant Canyon Field has an average daily production of over 7,000 barrels of oil from two wells, and has produced over 14 million barrels of oil since its discovery in 1983 (Petzet,

1991). These two fields combined have provided 48% of cumulative oil production for the state of Nevada. Both fields are water-driven, both are located on subsurface fault blocks that are truncated by unconformities, and both were drilled based on seismic information

(Veal, et al., 1988).

Exploration

Petroleum exploration in the Basin and Range is risky because of the complex geologic history of the region (Poole and Claypool, 1984; Cook, 1988). The two most important controls on reservoir development are 1) timing of hydrocarbon maturation and migration, and 2) spatial relation of source rocks to reservoir rocks (Poole and Claypool,

1984; Cook, 1988). The complex tectonics and high regional heat flow of the region have caused many accumulations to become overmature (Eaton, 1979; Cook, 1988). Poole and

Claypool (1984) stated that “ distribution of oil reservoirs is directly related to the presence of petroleum source beds.” Meissner (1978) identified two phases of hydrocarbon genera­ T-4010 81 tion in the Basin and Range region. The first episode occurred in the Late Paleozoic to Early

Mesozoic in response to burial. The second episode occurred in the Tertiary as a result of burial associated with Basin and Range deformation. Paleozoic source rocks, such as the

Ordovician Vinini Shale, Devonian-Mississippian Pilot Shale, and Mississippian Chainman

Shale, were subjected to both episodes, whereas Cretaceous and Tertiary source rocks, such as the Sheep Pass Formation, were subjected only to the second episode. Meissner (1978) suggested that no Paleozoic traps remain; rather, these early accumulations were either destroyed or remigrated to younger traps during Tertiary deformation.

Most of the existing Nevada oil fields were drilled based on structural mapping, and generally have a combination of structural and stratigraphic traps. In order to develop a successful exploration program in the Great Basin it is necessary to determine both structural and depositional history as well as timing of hydrocarbon maturation and migration for each individual basin (Poole and Claypool, 1984). Most of Lincoln County, Nevada, is cited as having low petroleum potential, with the extreme northern and southeastern portions of the county having moderate potential (Garside, et al., 1988; Figure 7.1). One well, the Amoco

Dutch John Unit #1, drilled in Lake Valley in northern Lincoln County, reported an oil show in Mississippian strata (Garside, et al., 1988). This is the only confirmed presence of hydrocarbons in Lincoln County reported to date.

The synthetic gamma ray log constructed for the Guilmette in the Pahranagat Range

(Plate 1) was plotted in gamma counts per second rather than API units, and so is not directly comparable to downhole gamma logs. However, trends were observed that may be of use to further exploration in Nevada. The Lower Member of the Guilmette generally exhibits high gamma ray values, while the Middle Member exhibits relatively low gamma ray values.

Additionally, the top of the Middle Member, which was interpreted as an unconformity, is T-4010

A*./88S&L ■ S ■ S ; S -v v

■lilip ■ S S r

• Oil fields

■ High potential

Moderate potential

i l l Low potential

□ Very low potential

FIGURE 7.1: Petroleum potential of Nevada (after Garside, et al., 1988). T-4010 83 marked by a sharp drop on the gamma ray curve. If this unconformity is regional, this spike on the gamma ray curve could be a marker on wireline logs of the Guilmette. The gamma ray curve for the Upper Member of the Guilmette is highly variable due to the variable lithology and was not interpreted to represent any regional trends.

The Middle Member is interpreted to have been cemented soon after deposition, relative to the rest of the Guilmette, and so could form a regional seismic marker. The thickness of this Member could be resolved at the seismic scale, depending on the lithology of overlying rocks.

Reservoir

The Guilmette reservoir at Grant Canyon Field is highly fractured, medium to dark gray dolomite, containing abundant Amphipora, hemispherical stromatoporoids, and bra- chiopod and crinoid fragments. The porosity is primarily fracture and fracture-connected vugs (Read and Zogg, 1988).

The best potential reservoir in the study area is thought to be the Middle Member, or the equivalent highstand-systems tract. The Guilmette in this area has porosity estimated at less than 1 %, while the Middle Member has fracture, interparticle, and intraparticle porosity estimated at up to 10%. Fracturing in the subsurface could significantly enhance this porosity. Subsurface fluids could also potentially enhance porosity through dissolution of carbonate material. Where present, the sandstones of the Upper Member (shelf margin wedge) are secondary targets as potential reservoirs. These units had very low (less than 1 %) observed porosity, but their carbonate cements make them excellent candidates for second­ ary porosity development in the subsurface. T-4010 84

CHAPTER 8 SUMMARY AND CONCLUSIONS

1. In this report, the Guilmette Formation was subdivided into three members based on lithology and depositional processes. These members are referred to as Lower, Middle, and Upper, and are separated by sharp, erosional contacts.

2. The Lower Member is composed of interbedded limestone and dolomite deposited as shallowing-upward cycles in a shallow shelf environment. Overall, this

Member appears to be transgressive. The base of this Member coincides with the base of the Guilmette Formation, and is placed at the base of the first-occurring, yellow-weathering dolomite. This boundary is approximately 80 feet above the Stringocephalus zone, which was included in the Guilmette by previous workers. For this study, the Stringocephalus zone is placed within the Fox Mountain Member of the underlying Simonson Dolomite.

3. The Middle Member forms a near-vertical cliff that is 130 to 175 feet thick. The base of this Member is erosional into the underlying Lower Member, and locally this surface is underlain by an algal-coated lag. This enigmatic Member exhibits characteristics of submarine debris flows, and was deposited in relatively deep water based on both petro- graphic and sedimentologic interpretations. The unit appears bedded in the vicinity of

Section 1 and at Section 2 with bedsets separated by recessive surfaces that may be interpreted as cycle tops. However, this bedding is anomalous when compared to the Middle

Member in other areas. The upper five to 10 feet of the Middle Member is composed T-4010 85 dominantly of grains that are generally less than 1 inch in longest dimension, and which diminish upward to form a graded bed. The thickness and grain size of this interval is locally variable due to the erosional unconformity at the upper contact of the Member, which is interpreted to represent a fall in relative sea level.

4 . The Upper Member is composed of interbedded limestone, dolomite, and sandstone deposited in subtidal to supratidal environments. These units appear to be somewhat cyclic, although not as well defined as the cycles in the Lower Member.

Sandstones may have been sourced by a local tectonic uplift. This Member generally appears to have an overall regressive trend.

5. In terms of sequence stratigraphy, the Lower Member is interpreted as the transgressive systems tract, the Middle Member is interpreted as a highstand systems tract, the top of the Middle Member is interpreted as a type 2 sequence boundary, and the Upper

Member is interpreted as a shelf margin wedge. Fischer plots were used as an objective means of supporting this interpretation.

6. Synthetic gamma ray logs plotted for the Guilmette using field data reveal trends that may be useful for regional correlations. In general, the Lower Member exhibits a high gamma curve, while the Middle Member exhibits a low gamma curve. The unconformity at the top of the Middle Member is represented by a drop on the gamma curve. This spike could be useful as a regional marker on downhole gamma logs of the Guilmette. T-4010 86

7. Because the Middle Member is interpreted to have been cemented relatively soon after deposition, it may be a regional seismic marker. The thickness of this Member may make it resolvable at the seismic scale. T-4010 8 7

REFERENCES CITED

Ackman, B.W., 1991, Stratigraphy of the Guilmette Formation, Worthington Mountains

and Schell Creek Range, southeastern Nevada: Master's thesis T-4012, Colorado

School of Mines, 207 p.

Armstrong, R.L., 1968, Sevier orogenic belt in Nevada and Utah: GSA Bulletin, v. 79,

p. 429-458.

Badiozamani, K., 1973, The Dorag dolomitization model-application to the Middle Ordo­

vician of Wisconsin: Journal of Sedimentary Petrology, v. 43, p. 965-984.

Biller, E.J., 1976, Stratigraphy and petroleum possibilities of lower Upper Devonian

(Frasnian and Lower Fammenian) strata, southwestern Utah: Master's thesis T-

1787, Colorado School of Mines, 105 p.

Boucot, A.J., J.G. Johnson, and W. Struve, 1966, Stringocephalus, ontogeny and distri­

bution: Journal of Paleontology, v. 40, p. 169-171.

Brett, C.E., and G.C. Baird, 1986, Symmetrical and upward-shallowing cycles in the

Middle Devonian of New York State and their implications for the punctuated

aggradational cycle hypothesis: Paleoceanography, v. 1, p. 431-445.

Burchfiel, B.C., and G.A. Davis, 1972, Structural framework and evolution of the

southern part of the Cordilleran Orogen, western United States: American Journal

of Science, v. 272, p. 97-118.

Burchfiel, B.C., and L.H. Royden, 1991, Antler orogeny: a Mediterranean-type orogeny:

Geology, v. 19, p.66-69.

Chamberlain, A.K., in prep., Stratigraphy and structure of the Timpahute Range, Nevada:

PhD dissertation T-4159, Colorado School of Mines

Cook, H.E., 1988, Overview: Geologic history and carbonate petroleum reservoirs of the

Basin and Range Province, western United States, in S.M. Goolsby and M.W. T-4010 g 8

Longman, eds., Occurrence and petrophysical properties of carbonate reservoirs in

the Rocky Mountain Region: RMAG Carbonate Symposium, p. 213-227.

Dickinson, W.R., 1977, Paleozoic plate tectonics and the evolution of the Cordilleran

continental margin, in J.H. Stewart, C.H. Stevens, and A.E. Fritsche, eds., Paleo­

zoic paleogeography of the western United States: SEPM Pacific Section, Pacific

Coast Paleogeography Symposium, p. 137-155.

Driese, S.G., and R.H. Dott, Jr., 1984, Model for sandstone-carbonate "cyclothems"

based on Upper Member of Morgan Formation (Middle Pennsylvanian) of northern

Utah and Colorado: AAPG Bulletin, v. 68, p. 574-597.

Dunham, J.B., and E.R. Olson, 1978, Diagenetic dolomite formation related to Paleozoic

paleogeography of the Cordilleran miogeocline in Nevada: Geology, v. 6, p. 556-

559.

______, and______, 1980, Shallow subsurface dolomitization of subtidally

deposited carbonate sediments in the Hanson Creek Formation (Ordovician-

Silurian) of central Nevada, in D.H. Zenger, J.B. Dunham, and R.L. Ethington,

eds., Concepts and models of dolomitization: SEPM Special Publication No. 28,

p. 139-161.

Dunham, R.J., 1962, Classification of carbonate rocks according to depositional texture, in

W.E. Ham, ed., Classification of carbonate rocks: AAPG Memoir #1, p. 108-121.

Dunn, M.J., 1979, Depositional history and paleoecology of an Upper Devonian

(Frasnian) bioherm, Mount Irish, Nevada: Master's thesis, SUNY at Binghamton,

133 p.

Eaton, G.P., 1979, Regional geophysics, Cenozoic tectonics, and geologic resources of

the Basin and Range province and adjoining regions, in G.W. Newman and H.D.

Goode, eds., Basin and Range Symposium and Great Basin Field Conference:

RMAG and Utah Geological Association, p. 11-39. T-4010 8 9

Embry, A.F., and J.E. Klovan, 1971, A Late Devonian reef tract on northeastern Banks

Island, Northwest Territories: Bulletin of Canadian Petroleum Geology, v. 19, p.

730-781.

Folk, R.L., and L.S. Land, 1975, Mg/Ca ratio and salinity: two controls over crystal­

lization of dolomite: AAPG Bulletin, v. 59, p. 60-68.

Garside, L.J., R.H. Hess, K.L. Fleming, and B.S. Weimer, eds., 1988, Oil and gas

developments in Nevada: Nevada Bureau of Mines and Geology Bulletin 104,

136 p.

Geological Society of America, 1983, Geologic time scale, 1 p.

Ginsburg, R.N., 1971, Landward movement of carbonate mud: new model for regressive

cycles in carbonates (abs.): AAPG Bulletin, v. 55, p. 340.

Goebel, K.A., 1991, Paleogeographic setting of Late Devonian to Early Mississippian

transition from passive to collisional margin, Ander Foreland, eastern Nevada and

western Utah, in J.D. Cooper and C.H. Stevens, eds., Paleozoic paleogeography

of the western United States-II: Pacific Section SEPM, vol. 67, p. 401-418.

Goodwin, P.W., E.J. Anderson, W.M. Goodman, and L.J. Saraka, 1986, Punctuated

aggradational cycles: implications for stratigraphic analysis: Paleoceanography,

v. 1, p. 417-429.

Grotzinger, J.P., 1986, Upward shallowing platform cycles: a response to 2.2 billion

years of low-amplitude, high frequency (Milankovitch band) sea level oscillations:

Paleoceanography, v. 1, p. 403-416.

Hamilton, W., 1987, Crustal extension in the Basin and Range Province, southwestern

United States, in M.P. Coward, J.F. Dewey, and P.L. Hancock, eds., Continental

Extensional Tectonics: Geological Society of London Special Publication No. 28,

p. 155-176.

Hintze, L.F., ed., 1985, Great Basin region, Correlation of Stratigraphic Units of North T-4010 9 0

America (COSUNA) Project: Tulsa, American Association of Petroleum

Geologists, 1 chart.

Hoggan, R.D., 1975, Paleoecology of the Guilmette Formation in eastern Nevada and

western Utah: Brigham Young University Geology Studies, v. 22, part 1, p. 141-

197.

Hurtubise, D.O., 1989, Stratigraphy and structure of the Seaman Range, Nevada, with an

emphasis on the Devonian system: PhD dissertation T-3688, Colorado School of

Mines, 443 p.

James, N.P., 1984, Shallowing-upward sequences in carbonates, in R.G.Walker, ed.,

Facies models (second edition): Geoscience Canada Reprint Series 1, p. 213-228.

Jayko, A.S., 1988, Late Cenozoic shallow crustal deformation in the Pahranagat and

adjacent ranges, southeastern Nevada (abs.): GSA Abstracts with Programs, v.

20, no. 3, p. 171.

Johnson, J.G., and C.A. Sandberg, 1988, Devonian eustatic events in the western United

States and their biostratigraphic responses: Devonian of the World, Volume 3:

Paleontology, paleoecology, and biostratigraphy: Canadian Society of Petroleum

Geologists Memoir 14, p. 171 -178.

Kellogg, H.E., 1963, Paleozoic stratigraphy of the southern Egan Range, Nevada: GSA

Bulletin, v. 74, p. 685-708.

Klein, G. deV., 1977, Clastic tidal facies: Champaign, Illinois, Continuing Education

Publication Company, 143 p.

Laporte, L.F., 1975, Carbonate tidal-flat deposits of the Early Devonian Manlius

Formation of New York State, in R.N. Ginsburg, ed., Tidal deposits: New York,

Springer-Verlag, p. 243-250.

Larsen, B.R., M.A. Chan, and S.R. Bereskin, 1988, Cyclic stratigraphy of the upper

member of the Guilmette Formation (Uppermost Givetian, Frasnian) west-central T-4010 9 1

Utah: Devonian of the World, Volume 2: Sedimentation: Canadian Society of

Petroleum Geologists Memoir 14, p. 569-579.

Levy, M., and N. Christie-Blick, 1989, Pre-Mesozoic palinspastic reconstruction of the

eastern Great Basin (western United States): Science, v. 245, p. 1454-1462.

Longman, M.W., 1982, Carbonate diagenesis as a control on stratigraphic traps: AAPG

Education Course Notes Series #21, 159 p.

Meissner, F.F., 1978, Petroleum geology of Great Basin, western Utah and Nevada

(abs.): AAPG Bulletin, v. 62, p. 888-889.

Muir, M.D., D.E. Lock, and C.C. von der Borch, 1980, The Coorong model for penecon-

temporaneous dolomite formation in the Middle Proterozoic McArthur Group,

Northern Territory, Australia, in D.H. Zenger, J.B. Dunham, and R.L. Ethington,

eds., Concepts and models of dolomitization: SEPM Special Publication No. 28,

p. 51-67.

Nilsen, T.H., and J.H. Stewart, 1980, The Antler orogeny--Mid-Paleozoic tectonism in

western North America: Geology, v. 8, p. 298-302.

Nolan, T.B., 1935, The Gold Hill mining district, Utah: USGS Professional Paper 177,

172 p.

Pettijohn, F.J., P.E. Potter, and R. Siever, 1987, Sand and sandstone (second edition):

New York, Springer-Verlag, 553 p.

Petzet, G.A., 1991, Water's arrival to prompt drilling in Nevada's Grant Canyon field: Oil

& Gas Journal, v. 89, no. 32, p. 90.

Phleger, F.B., and G.C. Ewing, 1962, Sedimentology and oceanography of coastal

lagoons in Baja California, Mexico: GSA Bulletin, v. 73, p. 145-182.

Poole, F.G., D.L. Baars, H. Drewes, P.T. Hayes, K.B. Ketner, E.D. McKee, C.

Teichert, and J.S. Williams, 1967, Devonian of the southwestern United States, in T-4010 9 2

D.H. Oswald, ed., International Symposium on the Devonian System, v. 1:

Alberta Society of Petroleum Geologists, p. 879-912.

Poole, F.G., and C.A. Sandberg, 1977, Mississippian paleogeography and tectonics of

the western United States, in J.H. Stewart, C.H. Stevens, and A.E. Fritsche, eds.,

Paleozoic paleogeography of the western United States: SEPM Pacific Section,

Pacific Coast Paleogeography Symposium, p. 67-85.

Poole, F.G., C.A. Sandberg, and A.J. Boucot, 1977, Silurian and Devonian paleogeo­

graphy of the western United States, in J.H. Stewart, C.H. Stevens, and A.E.

Fritsche, eds., Paleozoic paleogeography of the western United States: SEPM

Pacific Section, Pacific Coast Paleogeography Symposium, p. 39-65.

Poole, F.G., and G.E. Claypool, 1984, Petroleum source-rock potential and crude-oil

correlation in the Great Basin, in J. Woodward, F.F. Meissner, and J.L. Clayton,

eds., Hydrocarbon source rocks of the greater Rocky Mountain region: RMAG,

p. 179-207.

Read, D.L., and W. D. Zogg, 1988, Description and origin of the Devonian dolomite oil

reservoir, Grant Canyon Field, Nye County, Nevada, in S.M. Goolsby and M.W.

Longman, eds., Occurrence and petrophysical properties of carbonate reservoirs in

the Rocky Mountain Region: RMAG Carbonate Symposium, p. 229-240.

Read, J.F., 1975, Tidal flat facies in carbonate cycles, Pillara Formation,Western

Australia, in R.N. Ginsburg, ed., Tidal deposits: New York, Springer-Verlag, p.

251-256.

______, and R.K. Goldhammer, 1988, Use of Fischer plots to define third-order sea-

level curves in Ordovician peritidal cyclic carbonates, Appalachians: Geology, v.

16, p. 895-899.

Reso, A., 1960, The geology of the Pahranagat Range, Lincoln County, Nevada: PhD

dissertation, Rice University, 656 p. T-4010 9 3

______, 1963, Composite columnar section of exposed Paleozoic and Cenozoic rocks in

the Pahranagat Range, Lincoln County, Nevada: GSA Bulletin, v. 72, p. 901-918.

Sandberg, C.A., F.G. Poole, and J.G. Johnson, 1988, Upper Devonian of the western

United States: Devonian of the World, Volume 1: Regional Syntheses: Canadian

Society of Petroleum Geologists Memoir 14, p. 161-182.

Sarg, J.F., 1988, Carbonate sequence stratigraphy, in C.K. Wilgus, et al., eds., Sea-level

changes: an integrated approach: SEPM Special Publication No. 42, p. 155-181.

Shinn, E.A., 1973, Sedimentary accretion along the leeward, SE coast of Qatar Peninsula,

Persian Gulf, in B.H. Purser, ed., The Persian Gulf: Holocene carbonate sedimen­

tation and diagenesis in a shallow epicontinental sea: New York, Springer-Verlag,

p. 199-209.

Shinn, E.A., 1983, Tidal flat environment, in P.A. Scholle, D.G. Bebout, and C.H.

Moore, eds., Carbonate depositional environments: AAPG Memoir 33, p. 171-

210 .

Smith, S.M., 1984, Carbonate petrology and depositional environments of carbonate

buildups in the Devonian Guilmette Formation near White Horse Pass, Elko

County, Nevada: Brigham Young University Geology Studies, v. 31, part 1,

p. 117-139.

Stewart, J.H., 1971, Basin and Range structure: a system of horsts and grabens produced

by deep-seated extension: GSA Bulletin, v. 82, p. 1019-1044.

Tschanz, C.M., and E.H. Pampeyan, 1970, Geology and mineral deposits of Lincoln

County, Nevada: Nevada Bureau of Mines and Geology Bulletin 73, 188 p.

Tucker, M.E., 1985, Shallow-marine carbonate facies and facies models, in P.J. Benchley

and B.P.J. Williams, eds., Sedimentology: recent developments and applied

aspects: Geological Society of London Special Publication 17, p. 147-169.

Van Wagoner, J.C., H.W. Posamentier, R.M. Mitchum, P.R. Vail, J.F. Sarg, T.S. T-4010 9 4

Loutt, and J. Hardenbol, 1988, An overview of the fundamentals of sequence

stratigraphy and key definitions, in C.K. Wilgus, et al., eds., Sea-level

changes: an integrated approach: SEPM Special Publication No. 42, p. 39-45.

Veal, H.K., H.D. Duey, L.C. Bortz, and N.H. Foster, 1988, Grant Canyon and Bacon

Flat oil fields, Railroad Valley, Nye County, Nevada: Oil & Gas Journal, v. 86,

p. 67-70.

Warme, J.E., A.K. Chamberlain, and B.W. Ackman, 1991, The Alamo Event: Devonian

cataclysmic breccia in southeastern Nevada (abs.): GSA Abstracts with Programs,

v. 23, no. 2, p. 108.

Wernicke, B., 1981, Low-angle normal faults in the Basin and Range Province: nappe

tectonics in an extending orogen: Nature, v. 291, p. 645- 648.

Westgate, L.G., and A. Knopf, 1932, Geology and ore deposits of the Pioche District,

Nevada: USGS Professional Paper 171, 79 p.

Williams, W.L., 1984, Petrography and microfacies of the Devonian Guilmette Formation

in the Pequop Mountains, Elko County, Nevada: Brigham Young University

Geology Studies, v. 31, part 1, p. 167-186.

Wilson, J.L., 1975, Carbonate facies in geologic history: New York, Springer-Verlag,

471 p.

Witzke, B.J., and P.H. Heckel, 1988, Paleoclimatic indicators and inferred Devonian

paleolatitudes of Euramerica: Devonian of the World, Volume 1: Regional

Syntheses: Canadian Society of Petroleum Geologists Memoir 14, p. 49-63.

Yarmanto, in prep., Stratigraphy and depositional environments of a Devonian carbonate

breccia, northern Pahranagat Range, Nevada: Master's thesis T-4195, Colorado

School of Mines

Ziegler, W., and C.A. Sandberg, 1990, The Late Devonian standard conodont zonation:

Courier Forschungsinstitut Senckenberg, v. 121, p. 1-115.