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1988 Stratigraphy and sedimentology of Lower Cretaceous Sykes Mountain Formation, Bighorn Basin, Wyoming Hosny El-Desouky Ahmed Soliman Iowa State University
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Stratigraphy and sedimentology of Lower Cretaceous Sykes Mountain Formation, Bighorn Basin, Wyoming
Soliman, Hosny El-Desoiiky Ahmed, Ph.D.
Iowa State University, 1988
UMI SOON.ZeebRd. Ann Arbor, MI 48106
stratigraphy and sedimentology of Lower Cretaceous Sykes
Mountain Formation, Bighorn Basin, Wyoming
by
Hosny El-Desouky Ahmed Soliman
A Dissertation Submitted to the
Graduate Faculty in Partial Fulfillment of the
Requirements for the Degree of
DOCTOR OF PHILOSOPHY
Department: Earth Sciences Major: Geology
Signature was redacted for privacy. In Charg Work
Signature was redacted for privacy. r the Major Department
Signature was redacted for privacy. For the Graduate College
Iowa State University Ames, Iowa
1988 TABLE OF CONTENTS PAGE
INTRODUCTION 1 Purpose and Scope of Study 1 Field Work 3 Geologic Setting 6
CHAPTER I. STRATIGRAPHY 8 Previous Work 8 Stratigraphie Nomenclature 18 Stratigraphie Descriptions 22 Contacts 30 Basal contact 30 Upper contact 37 Regional Correlations 39 Summary 43
CHAPTER II. FACIES AND DEPOSITlONAL ENVIRONMENTS .... 46 Introduction 46 Description of Facies 48 Quartz Arenite Facies Association (QA) 49 Flat-Bedded and Parallel-Laminated Sandstone Facies (Sfl) 51 Planar Cross-stratified Sandstone Facies (Sp) ... 53 Trough Cross-stratified Sandstone Facies (St) ... 54 Hummocky Cross-stratified Sandstone Facies (She) 56 Convolute-Bedded Sandstone Facies (Scb) 58 Asymmetrical-Rippled Sandstone Facies (Sar) .... 59 Symmetrical-Rippled Sandstone Facies (Ssr) .... 61 Heterolithic Facies Association (SM) 63 Facies Sequence 65 Interpretation of the Depositional Environments ... 66 Comparison of Sedimentary Structures in Sykes Mountain Formation to Other Tidal Sediments .... 80 Summary 80
CHAPTER III. PALEOCURRENT ANALYSIS 85 Introduction - 85 Paleocurrent Measurement 86 Graphical presentation of data 91 Discussion 92 Summary 105
CHAPTER IV. PETROLOGY 107 Introduction 107 Petrology of Sandstones 108 Grain contacts 109 Cementation 110 i i i
Sources of secondary silica 110 Sources of carbonate cement Ill Origin of iron oxide cement 112 Heavy minerals 113 Provenance 117 Mudrocks 120 Dahllite concretions 122 Summary 125
CHAPTER V. TECTONICS AND SEDIMENTATION OF LOWER CRETACEOUS SEDIMENTARY ROCKS 127 Tectonics 127 Sedimentation 133 Summary 140
GENERAL SUMMARY AND CONCLUSIONS 142
BIBLIOGRAPHY 144
ACKNOWLEDGEMENTS 165 iv
LIST OF TABLES
PAGE
TABLE 1. Locations of measured sections 4
TABLE 2. Lithofacies characteristics and their probable depositional environments 50
TABLE 3. Sedimentary features in the Sykes Mountain Formation compared to other tidal sediments . . 81
TABLE 4. Summary of ripple measurements at different locations 89
TABLE 5. Summary of cross-stratification measurements 90
TABLE 6. Summary of current oriented fossils measurements 90
TABLE 7. Summary of all paleocurrent measurements at different locations 91 V
LIST OF FIGURES
PAGE
FIGURE 1. Location of Bighorn Basin, Wyoming (numbers refer to measured sections) 2
FIGURE 2. Development of nomenclature of the Lower Cretaceous rocks in the Bighorn Basin .... 19
FIGURE 3. Graphic section of Sykes Mountain Formation 29
FIGURE 4. Sykes Mountain Formation (SMF) above the Greybull Sandstone (GS) 32
FIGURE 5. Sykes Mountain Formation (SMF) above the Himes Member (HM) 33
FIGURE 6. Sharp erosional contact between Sykes Mountain and Cloverly formations 36
FIGURE 7. Thermopolis Shale (TS) above Sykes Mountain Formation (SMF) 38
FIGURE 8. Lithogenetic equivalents of the Sykes Mountain Formation and the adjacent formations 44
FIGURE 9. Flat bedding (FB) in the Sfl Facies underlying hummocky cross-stratification (HCS) in the She Facies 53
FIGURE 10. Large-scale planar cross-stratification (PCS) in the Sp Facies 55
FIGURE 11. Herringbone cross-stratification in the Sp Facies 56
FIGURE 12. Planar cross-stratified sandstone of the Sp Facies below the rippled sandstone of the Sar Facies 57
FIGURE 13. Sandwaves in the Sp Facies at Red Pryor Anticline 58 vi
FIGURE 14. Large-scale trough cross-stratification in the St Facies 59
FIGURE 15. Convolute bedding in the sandstones of the Scb Facies 60
FIGURE 16. Biogenic structures including burrows and trails on Facies Sar and Ssr 61
FIGURE 17. Straight-crested symmetrical ripples in the Ssr Facies 62
FIGURE 18. Lenticular bedding (LB) and wavy bedding (WB) in the SM Facies Association 65
FIGURE 19. Examples of vertical sequences in Sykes Mountain Formation 67
FIGURE 20. Alternation between the QA and SM facies associations 68
FIGURE 21. Rose diagrams from ripple measurements .... 93
FIGURE 22. Regional distribution of dominant and subsidiary paleocurrent directions from ripple measurements 94
FIGURE 23. Rose diagrams from cross-stratification measurements 95
FIGURE 24. Regional distribution of dominant and subsidiary paleocurrent directions from cross-stratification measurements 96
FIGURE 25. Rose diagrams from current-oriented pelecypod fossil measurements 97
FIGURE 26. Regional distribution of dominant and subsidiary paleocurrent directions from current-oriented pelecypod fossil measurements 98
FIGURE 27. Rose diagrams from measurements obtained from all indicators combined 99
FIGURE 28. Regional distribution of dominant and subsidiary paleocurrent directions from all indicators combined 100 vi i
FIGURE 29. Typical heavy mineral composition of the Sykes Mountain Formation (after Chisholm, 1963) 116
FIGURE 30. Distribution of the heavy mineral suites of Sykes Mountain Formation and Upper Lakota (after Chisholm, 1963) 118
FIGURE 31. Dahllite concretions from outside and inside 124
FIGURE 32. Map of portion of Wyoming province showing Laramide sedimentary basins and uplifts (modified from Dickinson et al., 1988) . . . 128
FIGURE 33. Major faults in the western overthrust belt (from Wiltschko and Dorr, 1983) .... 130
FIGURE 34. Diagrammatic map of the Wyoming foreland showing arcuate uplifts bounded by zones of thrusting or steep dip (thick curved lines) and deep basin areas (stippled), from Berg (1981) 134
FIGURE 35. Sedimentation on the overthrust belt and the adjacent foreland basin (modified from Wiltschko and Dorr, 1983) 139 1
INTRODUCTION
Purpose and Scope of Study
The present work is a detailed study of a sequence of clastic rocks which were deposited during the Early
Cretaceous in tidal flat and shallow water marine environments and which record the initial transgression of the Early Cretaceous Sea in the Western Interior. This sequence was defined as the "Sykes Mountain Formation" by
Moberly (1960), and was studied in detail by the writer along the margins of the Bighorn Basin, Wyoming (Figure 1).
The objectives of the present study include;
1. The measurement and description of critical
exposures and the establishment of local correlations.
2. The establishment of facies and facies associations.
3. Paleocurrent measurement and analysis.
4. Pétrographie studies of representative samples.
5. The interpretation of the environment of deposition.
This dissertation is divided into five chapters of which each chapter is written as a self-contained unit for future publication and can be read individually without reference to other chapters. Chapter I deals with the stratigraphy of the Sykes Mountain Formation in Bighorn
Basin and its equivalents in adjacent areas. Chapter II" 2
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FIGURE 1. Location of Bighorn Basin, Wyoming (numbers refer to measured sections) 3 concerns itself with facies and facies analysis and the interpretation of the environments of deposition. Chapter
III considers the measurement of paleocurrent indicators and their analysis, Chapter IV describes the petrology of the sandstones and the clay mineralogy of shales of the Sykes
Mountain Formation, and Chapter V discusses the tectonic and sedimentation history of the Early Cretaceous.
Field Work
Field work for this study was accomplished during the summers of 1984 and 1985 by measuring and describing exposures of the Sykes Mountain Formation at several critical localities using the Brunton Compass and Jacob's staff. Primary structures were studied and photographed.
Sandstone and shale samples were collected at the time of measurements.
The exposures measured and described were concentrated along the flanks of Sheep Mountain Anticline, Rose Dome,
Little Sheep Mountain Anticline, Sykes Mountain, Red Pryor
Anticline, and the Beaver Creek area on the eastern side of the Bighorn Basin (Figure 1). Other outcrops measured were located near the town of Cody on the western side of the basin, and near the towns of Thermopolis and Hyattville on the south-southeast margin of the basin (Figure 1 and Table 1). 4
TABLE 1. Locations of measured sections
No. Name Location
1 SMA. 1 NW. Sw. Section 13, T. 53 N. R. 94 W. Greybull North Quadrangle, Wyoming
2 SMA. 2 NW. NE. Section 14, T. 53 N. R. 94 W. Sheep Canyon Quadrangle, Wyoming
3 SMA. 3 Central Section 35, T. 53 N. R. 93 W. Greybull North Quadrangle, Wyoming
4 SMA. 4 SE. SW. Section 35, T. 53 N. R. 93 W. Greybull North Quadrangle, Wyoming
5 SMA. 5 SE. SE. Section 33, T. 54 N. R. 94 W. Sheep Canyon Quadrangle, Wyoming
6 SMA. 6 NW. NE. Section 29, T. 54 N. R. 94 W. Spence Quadrangle, Wyoming
7 SMA. 7 NW. SE. Section 18, T. 54 N. R. 94 W. Spence Quadrangle, Wyoming
8 LSMA. 8 NE. SW. Section 15, T. 55 N. R. 95 W. North Emblem Reservoir Quadrangle, Wyoming
9 LSMA. 9 South SE. Section 4, T. 55 N. R. 95 W. Lovell Lakes Quadrangle, Wyoming
10 LSMA. 10 SW. NW. Section 28, T. 56 N. R. 95 W. Lovell Lakes Quadrangle, Wyoming
11 RD. 11 SE. NW. Section 6, T. 54 N. R. 94 W. Spence Quadrangle, Wyoming
12 RD. 12 SE. NE. Section 1, T. 54 N. R. 95 W. Spence Quadrangle, Wyoming
13 RD. 13 NE. NE. Section 2, T. 54 N. R. 95 W. North Emblem Reservoir Quadrangle, Wyoming 5
TABLE 1. (Continued)
No. Name Location
14 SM. 14 SE. SW. Section 11, T. 57 N. R. 95 W. Sykes Spring Quadrangle, Wyoming
15 SM. 15 E. NE. Section 3, T. 57 N. R. 95 W. Sykes Spring Quadrangle, Wyoming
16 SM. 16 S. C. NW. Section 34, T. 58 N. R. 95 W. Sykes Spring Quadrangle, Wyoming
17 RPA. 17 C. NE. Section 32, T. 58 N. R. 96 W. Cowley Quadrangle, Wyoming-Montana
18 BCA. 18 Central NE. Section 1, T. 55 N. R. 92 W. Leavitt Reservoir Quadrangle, Wyoming
19 BCA. 19 C. N. SW. Section 24, T. 54 N. R. 92 W. Bear Creek Ranch Quadrangle, Wyoming
20 HY. 20 W. SE. Section 20, T. 49 N. R. 90 W. Weintz Draw Quadrangle, Wyoming
21 TH. 21 Highway 20 NW 1/4 Section 19, T. 43 N. R. 94 W. Hot Springs County, Wyoming
22 C. 22 NE. SW. Section 18, T. 52 N. R. 101 W. Devils Tooth Quadrangle, Wyoming
23 C. 23 NE Section 22, T. 51 N. R. 103 W. Devils Tooth Quadrangle, Wyoming
24 C. 24 NE. NW. Section 26, T. 52 N. R. 101 W. Cody Quadrangle, Wyoming 6
Geologic Setting
The Bighorn Basin lies within the Wyoming Province which is located within the Rocky Mountain Foreland between the Idaho-Wyoming Thrust Belt on the west and the North
American Craton to the east. The Wyoming Province is characterized by NW-SE trending mountain ranges, broad asymmetrical basins containing Paleozoic and Mesozoic sediments separated by Precambrian cored uplifts (Prucha et al., 1965; Grose, 1972; Kent, 1972; and Stearns, 1978).
Two major déformâtional events affected the Rocky
Mountain region during the period from Late Jurassic to
Eocene times. The first event was the Sevier Orogeny which took place from Late Jurassic to Paleocene time and produced a series of thrust faults in areas marginal to the North
American Cordilleran (Kent, 1972). The second déformâtional event was the Laramide Orogeny that overlapped the Sevier during the Late Cretaceous and extended to the Middle Eocene
(Love, 1960; Grose, 1972; Kent, 1972; and Bown, 1980). The
Laramide-style structural features in the Wyoming foreland include large asymmetric basins and arcuate basement thrusts
(Love, 1960; Berg, 1981; 1983; Blackstone, 1981; 1983).
However, an early episode of Laramide-style deformation occurred during the Early Cretaceous in Wyoming and Montana, as indicated by DeCelles (1984; 1986) and Kvale (1986), 7 indicating that the difference between the Sevier and
Laramide orogenies is one of style only.
Chamberlain (1945) and Prucha et (1965) reported the deposition of 2,400 to 4,500 meters of Paleozoic and
Mesozoic sediments in the Wyoming Province. According to
Bown (1980), in the Bighorn Basin, these deposits are overlain by 6,300 meters of Cenozoic strata which were derived from the adjacent uplifted areas. In the Rocky
Mountain region, the Cretaceous sediments were derived from mountain uplifts in the western Cordillera and the midcontinent region and were deposited in a broad asymmetrical subsiding trough (Kent, 1972). Richer (1960;
1962) reported that the Sykes Mountain Formation represents deposition from the initial transgression of the Early
Cretaceous (Middle Albian) sea which extended from the
Arctic Ocean southward toward the Gulf of Mexico. This transgressive cycle was followed by a series of transgressive-regressive cycles until the Middle
Maestrichtian (Kent, 1972). 8
CHAPTER I. STRATIGRAPHY
Previous Work
The early information about the geology and geography which existed in the American West before 1856 was provided by early travelers, hunters, and railroad surveys. In 1856,
Meek and Hayden published the results of their exploration of the Missouri and Platte Rivers. The first geologic maps of the Nebraska Territory resulted from these expeditions
(Meek and Hayden, 1856a; 1856b; 1858; 1861). Geologic observations in the Teton-Yellowstone area were provided by
Comstock (1873) and St. John (1883).
Eldridge (1894) was the first to refer to the Lower
Cretaceous rocks of the Bighorn Basin. The Dakota Formation was described by Eldridge (1894) as a commonly iron-stained quartzose sandstone of variable thickness that overlies
121.2 to 181.8 meters of Jurassic marine shales. Eldridge placed the beds that lie above his Dakota in the Colorado
Formation which he subdivided into the Benton and Niobrara formations.
The first detailed stratigraphical and structural studies of the Bighorn region were made by Darton (1906a).
He recognized Morrison sediments in northern Wyoming for the first time. The rocks which overlie the Morrison Formation 9 were given the name Cloverly Formation by Darton (1904). He correlated the Cloverly with the Lakota, Dakota, and Fuson formations of the Black Hills region (Darton, 1906a).
Darton (1906b) used the name Colorado Formation for the sediments above the Cloverly. He stated that: "The
Colorado Formation in the Bighorn region comprises the
Benton and Niobrara formations of the Rocky Mountain and
Great Plains region farther south and east .... The upper beds of the formation are gray shales, which probably represent the Niobrara Formation" (Darton, 1906b, p. 7).
In addition, Darton (1904, p. 399) stated that; "The basal member of the Benton consists of dark gray shales, in part sandy and of rusty brown color, with occasional thin beds of brown sandstone." This basal member is about 60.6 meters thick and is characterized by the presence of phosphate concretions at about 30.3 meters above its base. This member was called the rusty series (Darton, 1906a, p. 54).
Darton (1904, p. 400) gave the name "Mowrie beds" to the sediments overlying the rusty series. This sequence is divided into two units; the lower unit consists of 181.8 meters or more of black fissile shale containing carbonate and ironstone concretions, and a light-colored sandstone near its middle. The upper unit is composed of 45.5 to 91 meters of interbedded light gray, thin-bedded sandstone and 10
light gray, fish-scale-bearing shale. Later, this unit was referred to as the Mowry beds or the Mowry member (Darton,
1906a, p. 54; 1906b, p. 7).
Reports concerning the discovery of hydrocarbons in the
Bighorn Basin were first published by Washburne (1908). He referred to the thinly bedded sandstone and shale sequence at the base of the Colorado Formation as the "Rusty Beds" and interpreted it to represent the basal transgressive unit of the Cretaceous sea. Washburne (1908, p. 350) states that: "The 'Rusty Beds' are a constant feature of the base of the marine Cretaceous. Seemingly, they are as a group a true basal sandstone resting upon a rather smooth surface of erosion." He considered the base of the "Rusty Beds" to
represent the Lower-Upper Cretaceous boundary. He
attributed the disconformity at the base of the Cloverly sandstones below the "Rusty Beds" to the action of erosion
before the transgression of the Cretaceous sea. Washburne
(1909, p. 168) wrote: "... in all other places examined,
none of the Cloverly Formation above the basal sandstone is
certainly present, and in most places the entire formation
is absent. This hiatus is due to an unconformity separating the Cloverly Formation from the overlying marine Colorado
Shale." On the contrary, the Cloverly Formation is present
and underlies the Sykes Mountain Formation in all places 11
examined by the writer. Washburne (1908, p. 350) believed
that the gas obtained from the well drilled in 1907 at
Greybull was "either from the "Rusty Beds" at the base of
the Colorado, or from some sandstone in the shales less than
30.3 meters above the "Rusty Beds", with the probabilities strong in favor of the latter position."
Little was added to Barton's description of the sequence in the Bighorn Basin by Fisher (1906). However, he
recognized a gypsum bed in the upper part of the Morrison south of Cody (Fisher, 1906, p. 58), and some coal in the
basal Cloverly on No Wood Creek (Fisher, 1906, p. 53).
Hewett (1914) briefly described the Morrison and
Cloverly formations along the Shoshone River near Cody. He separated the Cloverly Formation from the Morrison Formation
and placed the contact between them at the top of the
highest maroon clays.
Hintze (1915) applied the term Greybull sandstone to
the lenticularly shaped sandstone body at the top of the
Cloverly Formation which he believed to be the gas-producing
horizon at Greybull Dome, at Greybull, Wyoming. According
to him, the Cloverly-"rusty beds" boundary was either a
disconformity or a slight angular unconformity in accordance
with Washburne's belief. 12
Lupton (1916) applied the term Thermopolis Shale to all of the strata between the Cloverly and the Mowry formations.
These strata had been previously assigned to the Benton
Formation (Darton, 1904) or the Colorado Formation (Darton,
1906a). However, Lupton (1916) subdivided the Colorado
Formation of Darton (1906a) into the Thermopolis, Mowry, and
Frontier formations of which the Thermopolis Formation included the "Rusty Beds" or Sykes Mountain Formation at its base. He reported that the Thermopolis Shale is conformable with the underlying Greybull sand of the uppermost Cloverly and with the overlying Mowry Shale.
Hewett and Lupton (1917, p. 18-20) described the
Morrison Formation and its contact relationship with the underlying Sundance Formation and the overlying Cloverly.
They formally established "the Greybull sandstone member of the Cloverly Formation" by assigning the 3 to 6 meters of yellowish-gray sandstone locally lying at the top of the
Cloverly Formation to the member. They named the persistent white sandstone near the middle of the Thermopolis Shale the
"Muddy Sand" as it was referred to by the drillers.
Ziegler (1917a; 1917b) suggested that the Morrison-
Cloverly contact should be placed at the top of the uppermost maroon clays. He considered the contact between the Cloverly and the Thermopolis to be disconformable in places where the basal 'Rusty Beds' lie directly on what he 13 thought were the variegated clays of the Morrison. He reported the occurrence of small phosphatic concretions near the base of the "Rusty Beds". According to him, the "Rusty
Beds" are the initial deposits of a transgressive sea while the Thermopolis and Mowry shales record quiet, shallow water marine conditions.
Hares (1917) assigned the name "Pryor Conglomerate" to the lowest member of the Cloverly Formation which is well exposed along the west flank of the Pryor Mountains.
Reeside (1923) described the small molluscan faunae from the middle shale of the Dakota Formation of north central
Colorado and southeastern Wyoming and regarded them to be of
Early Cretaceous age. Wilmarth (1925) assigned the "Rusty
Beds" to the Cloverly Formation. She described the lower
Cloverly as being "partly conglomeratic."
Lee (1923) introduced the term "Dakota Group" for the
Dakota Formation of northern Colorado and divided it into five members traceable northward to Greybull, Wyoming.
These were, from bottom to top, the Lower Sandstone, Lower
Shale, Middle Sandstone, Upper Shale, and Upper Sandstone.
Lee (1927) correlated his five members with the three members of the Cloverly and with the Thermopolis Shale, and the Muddy Sandstone; respectively. He reported that his
"Middle Sandstone" was equivalent to the "Greybull
Sandstone" of the Bighorn Basin. 14
There has been a controversy regarding the Morrison-
Cloverly and Cloverly-Thermopolis contacts. Hewett (1914) drew the Morrison-Cloverly contact at the top of the uppermost reddish beds and the Cloverly-Thermopolis contact below the dark gray shales of the Thermopolis. Hewett's
(1914) suggestion regarding the position of these contacts was followed by many subsequent workers (e.g., Pierce and
Andrews, 1940; Pierce, 1948).
Knappan and Moulton (1930) believed that the lower thinly bedded sandstones of the "Rusty Beds" exposed to the north of the Bighorn Basin were equivalent to the Greybull
Sandstone. As a result, they included them in the Cloverly
Formation, thereby increasing its thickness northward.
Pierce (1948) recognized the "Rusty Beds" on the western side of the Bighorn Basin, but he included them with the
Cloverly Formation on his map. Since the "Rusty Beds" and the Thermopolis Shale are stratigraphically equivalent to the Fall River Sandstone and the Skull Creek of the Black
Hills, respectively, Richards (1955, P. 41-42) considered the "Rusty Beds" to be the uppermost member of the Cloverly
Formation following the suggestion of earlier workers.
Mills (1956, p. 19), on the basis of a subsurface study reported that the Thermopolis Shale occurred over the
Greybull Sandstone and under the Muddy Sandstone. He 15
divided the Thermopolis Shale into "an upper shale, a middle silt, and a lower silt and sand." He correlated the middle silt unit with the "Dakota Silt" in Montana and applied the
name "Rusty Beds" to the lower silt and sand unit.
Working on the subsurface stratigraphy in the Wind
River Basin, Burk (1956, p. 30) stated that; "The Fall
River interval can be traced westward and northward into the
Thermopolis "Rusty Beds", which overlie the Greybull
Sandstone of the Bighorn Basin." Burk (1956, p. 31) placed
the "Rusty Beds" at the base of the Thermopolis Shale in the
western part of the Wind River Basin and correlated them
with the upper Inyan Kara Group in the eastern part of the
Basin. Faulkner (1956, p. 40) reported that the Thermopolis
Shale and the Skull Creek Shale are equivalent and
conformably overlie the Inyan Kara Group.
Wilson (1958, p. 78), in his section from the
northwestern Bighorn Basin to the northern Black Hills,
reported a change in thickness for the middle Cloverly
member without a specific pattern of thinning to the east.
He states: "The upper Cloverly sandstone to the west is
intimately associated with the 'Rusty Beds', the sequence
present representing clastics derived from the Bear River
complex to the west." 16
Waage (1958) discussed the stratigraphy of rocks at the base of the Cretaceous sequence in the Western Interior. He
(Waage, 1958, p. 76) states that: "The 'erosional surface' recognized by Washburne (1908) is the transgressive disconformity which lies at the base of the Fall River sandstone in the Black Hills region." In addition, he reported that he found the disconformity in the eastern part of the Bighorn Basin and along the eastern flank of the
Bighorn Mountains.
Moberly (1956; 1960; 1962) studied the Upper Jurassic
Morrison, Lower Cretaceous Cloverly and Sykes Mountain formations in the Bighorn Basin. He established the terminology currently in use in the basin. He proposed the name Sykes Mountain Formation for the "Rusty Beds" and the
Greybull Sandstone in 1960. He formally defined the
Cloverly Formation in the Bighorn Basin and divided it into three members: the Pryor Conglomerate, the Little Sheep
Mudstone, and the Himes. Moberly placed the Cloverly
Formation's lower contact at the base of the Pryor
Conglomerate or at the lowest occurrence of volcanic debris which is evidenced by the presence of "bentonites, partly bentonitic mudstones and claystones, highly siliceous rocks such as cherty siltstones, tuffaceous mudstones and spotted to white cherts, and variegated rocks especially colored 17 shades of pale-red, light dark-purple, pale greenish-yellow and neutral gray" (Moberly, 1960, P. 1145).
Richer (1958; 1960; 1962) studied the stratigraphy and paleontology of the Thermopolis Shale. He divided it into four members from base to top; the "Rusty Beds", lower black shale, thin silty shale, and upper black shale. Eicher
(1962, p. 75) states that "the Rusty Beds are marine or marginal marine sediments which overlie a regional surface of transgression formed by the initial advance of the Lower
Cretaceous sea into the Western Interior." He also reported that the "Rusty Beds" become sandier westward "from the east flank of the Bighorn Basin." The Rusty Beds on the western flank of the basin are sandier than those on the eastern flank.
Mirsky (1962a; 1962b) recognized four units that correspond to those recognized by Darton (1904). He considered the uppermost unit, a "sequence of interbedded black shales and thin sandstone" (rusty series), to be the lower part of the Thermopolis Shale.
Ostrom (1970) divided the Upper Jurassic-Lower
Cretaceous sequence into calcareous and noncalcareous units.
His calcareous units are equivalent to the Morrison
Formation of Moberly (1960) and the noncalcareous units correspond to the Cloverly Formation. His unit VIII corresponds to the lower part of the Sykes Mountain 18
Formation or the Greybull Sandstone of some other investigators. He assigned a Middle Albian age to the Sykes
Mountain Formation.
Studying the paleoenvironments and the tectonic significance of the Upper Jurassic Morrison and Lower
Cretaceous Cloverly formations, Kvale (1986) divided
Moberly's (1960) Himes Member into a lower unit including channel sandstones and overbank deposits and an upper unit
(referred to as the Greybull Sandstone) of fluvial and estuarine origin.
Several unpublished reports by Iowa State University students dealing with the general geology of different areas in the Bighorn Basin have been completed and are on file in the Iowa State University Library. These include Kozimko
(1977), Ladd (1979), Manahl (1981), Reppe (1981), Dunn
(1986), Kvale (1986), Noggle (1986), and Mantzios (1986).
Figure 2 is a chart showing the development of nomenclature for the Lower Cretaceous rocks in the Bighorn Basin.
Stratigraphie Nomenclature
The term "Sykes Mountain Formation" was defined by
Moberly (1960) for those Lower Cretaceous sediments of a distinctive rusty color exposed near Sykes Mountain and that overlie the Cloverly Formation and underlie the Thermopolis WASHBURNi HIUTZE UIPTOH HEWETT ZIEGLER DAHTON KNAPPAN PIERCE MILLS EICHER HOBERLY ossaoH KVALE THIS and (1906) (1908) (1915) (1916) (1917) and (1948) (1956) (I960, (1970) (1986) REPORT LUPTON HOULTON 1962) (1917) (1950)
Howry Mowry Howry Howry Howry Mowry Howry Mowry Shale Howry Mowry Howry Mowry Shale Shale Shale Shale Shale Shale Hbr. Shale Shale Shale Hbr. Hbr. sk^U
gUpper asThermO' Shell Shell Spolis Creek Creek
1 Muddy Muddy Muddy Muddy Muddy Muddy Muddy Middle Ss. Ss. Ss. Ss. 1 3a.Hbr Ss. Ss. 1 Upper s Shale i m S S Dakota Middle OM Silt §
"Rusty Lower ii Series" Shale 5
SYKES SYKES Rusty 'Rusty CLOVER- "Rusty Rusty SYKES Rusty Rusty HTN. PM. MTN. PM, Beds Beds Beds Beds" LY ÏM. Beds" Beds MTH. PH. Greybull Ss. Greybull Ss Greybull Ss. Greybull Ss. VIII Greybull Ss. Greybull VII Ss. Himes VI lower Hbr. Himes §i §1 Little CO B* Lit lie i ii Sheep Sheep ii il ii il i Mds. Mds. S3 Pryor 11 §8 §8 §1 ii il ii i IV Pryor 1 Congl Congl.
FIGURE 2. Development of nomenclature of the Lower Cretaceous rocks in the Bighorn Basin . 20
Shale. Regarding the Sykes Mountain Formation, Moberly
(1960, p. 1149) states that; "This formation includes the
'Rusty Beds', 'Dakota Silt', 'Dakota Sandstones', and
'Greybull Member' of the Cloverly Formation, of various
authors." Before Moberly's naming of the Sykes Mountain
Formation in 1960, these strata were assigned different
names by different authors. It is the basal member of
Darton's (1904) Benton Formation and what he called the
Rusty Series (Barton, 1906a, p. 54). He reported a
thickness of about 60.6 meters and the presence of phosphate
concretions at about 30.3 meters above its base. Washburne
(1908) described it as a thinly bedded sandstone and shale sequence at the base of the Colorado Formation and called it
the 'Rusty Beds'. He interpreted it to be the basal unit
deposited by the transgressive Cretaceous sea. Hintze
(1915) called the basal part of the Lower Benton Shale the
"Rusty Beds" and noted that it was overlain by black shales.
Ziegler (1917a) placed the "Rusty Beds" at the base of the
Thermopolis Shale and reported the occurrence of phosphate
concretions near its base. Wilmarth (1925) placed the
"Rusty Beds" above the Greybull Sandstone member of the
Cloverly Formation. Knappan and Moulton (1930) thought that
the lower part of the "Rusty Beds" was equivalent to the
Greybull Sandstone and included it in the Cloverly 21
Formation. Pierce (1948) also included the "Rusty Beds" in the Cloverly. Mills (1956) assigned the name "Rusty Beds" to the lower silt and sand unit of his Thermopolis Shale.
Burk (1956) placed the "Rusty Beds" at the base of the
Thermopolis Shale in the western part of the Wind River
Basin and reported that the Fall River Formation could be
traced laterally into the Thermopolis "Rusty Beds". Eicher
(1960; 1962) included the "Rusty Beds" in the Thermopolis
Shale as its lowest member because the marine sediments of
the "Rusty Beds" are more closely related to the marine
black Thermopolis Shale above than to the nonmarine light-
colored sediments of the Cloverly Formation below.
Moberly (1956) included the Greybull Sandstone of many
earlier workers with the "Rusty Beds" in his Crooked Creek
Formation. He later (1960) renamed this unit the Sykes
Mountain Formation. Moberly (1960, p. 1147) stated:
"Perhaps the Sykes Mountain should be designated only a
basal transgressive member of the Thermopolis Shale, but its
distinct lithology and differing origin argue against that
classification."
In the present report, the writer considers the Sykes
Mountain Formation to be that sequence of thin, interbedded,
rusty-brown, fine-grained sandstones, siltstones, and dark
gray shales which overlie the nonmarine Cloverly Formation 22
and grade upward into the overlying marine Thermopolis
Shale.
Stratigraphie Descriptions
The term Sykes Mountain Formation as used in this
report includes the Dakota Sandstone of Eldridge (1894), the
Rusty Series of Darton (1906a), the Rusty Beds of Washburne
(1908), Hintze (1915), Hewett and Lupton (1917), Mills
(1956), and Eicher (1960). It comprises that sedimentary sequence overlying the nonmarine Cloverly Formation and underlying the marine Thermopolis Shale. Moberly (1960) measured and described the Sykes Mountain Formation at its type section "northwest of Sykes Mountain between Crooked
Creek and Gypsum Creek about 1 mile south of the Montana
State line, in the northeastern quarter of sec. 25, T. 58
N., R. 96 W." There the sandstones, siltstones, ironstones, and shales that constitute the unit attain a thickness of
41.2 meters (Moberly, 1960). The formation varies in thickness from about 30.3 to about 90.9 meters (Moberly,
1962).
Burk (1957) recorded a thickness of 30.3 to 36.4 meters for the "Rusty Beds", which according to Eicher (1958; 1960; and 1962), represent the basal member of the Thermopolis
Formation and consist of 36.4 meters of tan and rusty. 23 thinly bedded and interlaminated sandstone, siltstone, and dark gray shale. Kozimko (1977) and Reppe (1981) reported that the average thickness of Sykes Mountain Formation is
30.3 meters in the Greybull North and Devils Kitchen
Quadrangles. Ladd (1979) reported a thickness of 30.3 meters of the Sykes Mountain Formation in Sheep Canyon
Quadrangle, Wyoming. Manahl (1981) measured a thickness of
48.5 meters where he described the Sykes Mountain Formation and pointed out that the variation in thickness for the
Sykes Mountain Formation and Thermopolis Shale was due to their interfingering. Dunn (1986) recorded a thickness of
43 meters for the Sykes Mountain Formation near the Beaver
Creek ranch, and he also reported a variation in thickness due to the interfingering of the Sykes Mountain Formation with the Thermopolis Shale. Noggle (1986) measured about
36.4 meters of rusty brown siltstones and dark shales of the
Sykes Mountain Formation exposed in the Leavitt Reservoir
Quadrangle.
Although the Sykes Mountain Formation is commonly poorly exposed, there are several excellent outcrops in some areas within the Bighorn Basin. These areas include Sheep
Mountain Anticline, Rose Dome (north east of the North Nose of the Sheep Mountain Anticline), Little Sheep Mountain
Anticline, Sykes Mountain, Red Pryor Anticline, the Beaver 24
Creek area, the vicinity of Hyattville, the area north of the town of Thermopolis, and southwest and northwest of the town of Cody. In these areas, the maximum thickness for the
Sykes Mountain Formation was 80.5 meters measured at
Thermopolis and the least was 21.1 meters measured near
Hyattville. In other locations, the thickness varies from
30 to 50 meters. This variation in thickness is due to the interfingering of the Sykes Mountain Formation with the overlying black shale of the Thermopolis. In addition, some exposures were not completely measured because the upper parts of the formation were covered.
Lithologically, the Sykes Mountain Formation consists of thin, interbedded and interlaminated, iron-oxide stained sandstones, siltstones, and gray shales. The sandstones are fine- to very fine-grained quartz arenites and dominate the lower part of the formation. Sandstone bodies near the base of the formation are thicker and coarser grained than those in the upper parts. They become thinner and finer-grained upwards. They are stained and cemented by hydrated iron- oxide and contain ironstone concretions and iron-rich nodules. Spherical and elliptical pyrite nodules occur in some sandstones and siltstones. The sandstone bodies are tabular, sheet- or blanket-shaped, flat bottomed, and sometimes erosive, laterally persistent and traceable for 25 miles. Some areas are characterized by the presence of sandstone channels near the base and others are not.
Sandstone beds vary in thickness from around one to two meters near the base of the formation (rarely exceeding two meters) to a few centimeters in the upper part. They are well indurated, both ferruginous and calcareous becoming more calcareous upwards. They are pale yellow to dark yellowish-orange, rusty, and dark reddish-brown. Thin ironstone beds that also occur near the base can be traced for long distances.
A variety of primary structures occur in these sandstones, including bedding, lamination, cross- stratification, herringbone cross-stratification, reactivation surfaces, convolute bedding, and symmetrical and asymmetrical ripples. Scour and fill features occur.
Biogenic structures including burrows and trails of organisms are abundant and may disturb the above-mentioned structures. Most of these structures are similar to those found in Recent tidal flat environments, although they may occur in many other environments. Near the base of the formation at Rose Dome (Section RD. 12, Table 1), there is a
90 cm thick, dark brown sandstone channel which has an ironstone band at its base, laminated sandstone and siltstone in the middle, and coquinoidal sandstone at the 26 top containing a high content of shell fragments and exhibiting hummocky cross-stratification indicating storm action. This sand unit changes laterally in thickness and in composition from a coquinoidal sandstone to a quartz arenite and pinches out laterally over a short distance. At some locations, fossil pelecypod casts occur on the surface of a few sandstone bodies near the base of the Sykes
Mountain Formation. Cone-in-cone structures occur at the base of some mudstone units. The Sykes Mountain Formation overlies either the brown or red sandstones, siltstones and mudstones of the Himes Member of the Cloverly Formation or the white quartz arenite of the Greybull sandstone.
Wherever the sandstones of the Sykes Mountain Formation overlie the Greybull sandstones, they can be distinguished on the basis of the following criteria reported by Moberly
(1962), Ladd (1979), Manahl (1981), and the writer's field observations;
1) Sykes Mountain sandstone bodies are tabular, but the
Greybull sandstone bodies have a lenticular geometry.
2) Sykes Mountain sandstones are laterally persistent and
traceable over long distances, while the Greybull
sandstones pinch out laterally over short distances.
3) Herringbone cross-stratification occurs in Sykes
Mountain but does not occur in the Greybull sandstones.
4) Sykes Mountain sandstones possess smaller scale cross- 27
stratification than Greybull sandstones.
5) Iron oxide nodules and concretions are abundant in the
Sykes Mountain Formation but not in the Greybull
sandstones.
Going up-section, above every sandstone body, the sequence consists- of intercalations of yellow, very fine grained quartz arenite and siltstone and gray shale. In the lower part of the formation, the sandstone and siltstone dominate, whereas in the upper part shale dominates. In the upper parts, the shale is more fissile and darker in color than in the lower parts. These sediments are characterized by the presence of flaser bedding, lenticular bedding and wavy bedding and ripple marks on the top of the sand lenses which are usually bioturbated. Wood fragments, fossil leaves, and black carbonaceous materials and iron oxide pellets occur. The uppermost part, which grades upwards to the marine black Thermopolis Shale, is dominated by dark gray fissile shale. The Sykes Mountain-Thermopolis Shale contact is characterized by the presence of phosphate concretions which occur in every measured section throughout the Bighorn Basin. The formation as a whole represents an overall fining upward sequence which is composed of several small sequences. The dark gray fissile shale units thicken as the sandstones thin. The Sykes Mountain Formation as a 28 whole changes lithologically along the eastern flank of the
Bighorn Basin from west to east. The sediments in the
Beaver Creek area are much finer-grained than those in the
Sheep Mountain Anticline area. In addition, the sequence becomes more sandy in the western and southern parts of the
Bighorn Basin.
A few casts of poorly preserved pelecypod fossils which may be nonmarine occur near the base of the Sykes Mountain
Formation in many localities throughout the basin. Kvale
(1986) identified two species of pelecypods; Unio
(Lamps il is) firri and Unio (Elliptic) douqlassi. Eicher
(1962) reported that he found casts of pelecypods and gastropods at one locality in the lowermost portion of his
Rusty Beds. He also reported the abundance of fossil leaves, wood fragments, silicified wood, trails, and burrows especially in the sandstone units. Eicher (1962) reported that no foraminifera were found, however, "one tiny marine protozoan, a tintinnid about 0.10 mm long, nearly always flattened in preservation, is common."
The Sykes Mountain Formation weathers to a rusty-brown or yellow color. It is often iron-oxide stained, and contains abundant ironstone concretions and laminae. Figure
3 is a typical section of Sykes Mountain Formation. 29
LOCATION: gasi-^ac E. NE. SEC. 3 R.95W. T. 57 N. SYKES SPRING QUAD. WYOMING
KEY:
Very Fine Quartz Arenite Fissile Shale Fe-Bond
Plane Bedding and Lamination Planar Cross- Stratification . VV. L/ Trough Cross- u/f.V.w Stratification Ripple Marks
SkfBiPSStX.nC'*':" TE^2gg::r2i2^^ Flaser Bedding Wavy Bedding rnwml CO i##m (SCGJ Lenticular Bedding Bioturbation C/3 Dahllite Concretions
m?#*:
#m
FIGURE 3. Graphic section of Sykes Mountain Formation 30
Contacts
Basal contact
The contact between the Sykes Mountain and Cloverly formations is sometimes difficult to establish, especially where the tabular sandstones of the Sykes Mountain Formation directly overlie channel sandstones of the Himes Member of the Cloverly Formation, because of the similar lithology of the sandstones below and above. However, it is possible to select the contact between the two formations by careful tracing of the thinly bedded sheet sandstones of the Sykes
Mountain Formation laterally from an adjacent area past that of the channel edge of the Cloverly Formation. In addition, the light-gray, thinly bedded, carbonaceous, fissile, silty shales and shaly siltstones so characteristic of the Sykes
Mountain Formation do not occur in the Cloverly Formation.
Eicher (1958; 1960; 1962) stresses the importance of the occurrence of spherulitic siderite in other Cretaceous strata in the western interior and uses them to establish the contact. He reported that at various locations in the
Western Interior where Lower Cretaceous transgressive marine sediments overlie nonmarine Lower Cretaceous deposits, small, orange, iron-oxide spherulites occur in the claystones directly below the contact. Waage (1955; 1958; 31
1959a; 1959b) reported a similar occurrence of siderite spherulites just below the base of the transgressive Fall
River sandstone and the contact separating the Lower
Cretaceous nonmarine and marine sediments in the Black Hills and Colorado region. He correlated this siderite-rich zone with the initial marine transgression. These spherulites may have formed under similar environmental conditions just prior to the Early Cretaceous transgression and mark a regional contact overlain by sediments deposited during or after the transgression of the sea (Eicher, 1960). In many places within the Bighorn Basin, the basal units of the
Sykes Mountain Formation contain these siderite spherulites.
The basal contact of the Sykes Mountain Formation is placed below the rusty-brown, well-indurated, fine-grained, thinly bedded, tabular quartz arenite or dark gray mudstones with siderite nodules and black carbonaceous plant debris, and above the commonly discontinuous, lenticular, light gray and white, noncalcareous, friable, silty, fine-grained quartz arenites of the Greybull Sandstone (Figure 4), or above the red and brown blocky mudstones, and siltstones of the Himes Member (Figure 5).
There are different opinions regarding the nature of the contact between the Cloverly Formation and Sykes
Mountain Formation. Washburne (1909) reported a hiatus due 32
FIGURE 4. Sykes Mountain Formation (SMF) above the Greybull Sandstone (GS)
to an unconformable relationship between the Cloverly
Formation below and the Rusty Beds above. He interpreted this hiatus to be the result of erosion prior to the encroachment of the Early Cretaceous sea. Hintze (1915) reported a disconformity or slight angular unconformity between the two formations. Lupton (1916) suggested a conformable relationship between the Thermopolis Shale and the Greybull sand. Ziegler (1917a) considered the contact to be disconformable where the Rusty Beds lie directly over the "Morrison". Faulkner (1956) considered the contact of the Thermopolis Shale and its equivalent "Skull Creek" with 33
FIGURE 5. Sykes Mountain Formation (SMF) above the Himes Member (HM)
the underlying Inyan Kara Group to be conformable. Waage
(1958) discussed the erosional surface (disconformity) recognized by Washburne (1908) in the Black Hills region and reported that he found a similar situation in the Bighorn
Basin and along the Bighorn Mountains. Eicher (1960) included the Rusty Beds and the Greybull sandstone in the
Thermopolis Shale. He considered the relationship between
Thermopolis Shale and Cloverly Formation to be disconformable in spite of his statement (p. 20); "the precise contact between the Rusty Beds and the Cloverly is 34 difficult to select consistently in the Bighorn Basin."
According to Moberly (1960, p. 1147), the sediments of the
Sykes Mountain Formation "lie with very sharp contact, representing a hiatus, on the nonmarine variegated claystones and channel sandstones of the Himes Member of the
Cloverly Formation." Moberly considered the Sykes Mountain
Formation to rest disconformably on the Cloverly Formation.
The character of parallel stratification and the high fissility of the Sykes Mountain Formation deposits in contrast to the blocky and variegated mudstones of the
Cloverly Formation have been used by Ostrom (1970) as criteria for placing the contact between Cloverly and Sykes
Mountain Formation. Ostrom did not find any evidence of an erosion surface between the Cloverly Formation and the
Thermopolis Shale. He suggested that the local absence of his unit IV of the Cloverly Formation was due to nondeposition and not to erosion. Moreover, he established a Late Aptian to Early Albian age for the Cloverly Formation and a Middle Albian age for the overlying Sykes Mountain
Formation.
Weimer (1983; 1984), studying the tectonics and sea level changes and their effect on the deposition and distribution of unconformities within the Cretaceous of the
Western Interior, identified "nine major regional to near 35
regional unconformities". One of these major unconformities
has been estimated as late Aptian-early Albian,
approximately 100 m. y. before the present (Weimer, 1983;
1984). This corresponds with the time following the
deposition of the nonmarine Cloverly Formation and preceding
the deposition of the Sykes Mountain Formation and the
Thermopolis Shale (Middle Albian). Kvale (1986) considered
the Sykes Mountain Formation to conformably overlie the
Cloverly Formation.
Generally, there are two different points of view
regarding the contact relationship of the Cloverly-Sykes
Mountain formations; the first opinion considers the two
formations as being conformable while the latter considers
them to be disconformable. The writer favors the second
opinion since he has recognized a very sharp erosional
contact between the two formations at some outcrops in the
area of study (Figure 6). In addition, this opinion is supported by the major unconformity identified by Weimer
(1983; 1984), as well as by the absence of the Cloverly
deposits in some places, either due to erosion or lack of
deposition as mentioned above. Moreover, the writer
believes that the development of paleosols (Mantzios, 1986)
on the Cloverly Formation took place during that period of
erosion prior to the deposition of the Sykes Mountain 36
Formation. The transgression of the Sykes Mountain sea probably resulted in the alteration of the soils and their pedogenic features (Mantzios, 1986). However, although the writer believes that the relationship between the Sykes
Mountain and the Cloverly formations is unconformable, a conformable relationship may occur in areas other than those checked by him since "unconformity is used for a major break that normally, but not always, can be traced over large areas" (Weimer, 1984, p. 8).
FIGURE 6. Sharp erosional contact between Sykes Mountain and Cloverly formations 37
Upper contact
The Sykes Mountain Formation-Thermopolis Shale contact is gradational and conformable (Figure 7). The gradation of the Sykes Mountain Formation upward to the Thermopolis Shale and the repetition of Sykes Mountain lithologies upward in the section (in the Middle unit of the Thermopolis) make it difficult to establish the contact. Moberly (1956; 1960) placed the contact at the top of the uppermost sandstones, siltstones, and ironstones of paralic origin. Eicher (1960, p. 19) placed the upper contact of the Rusty Beds (Sykes
Mountain Formation) above a "flaggy siltstone unit below a
9.1 meter shale containing a bed of dahllite concretions."
Ostrom (1970, p. 19) did not recognize any consistent contact or specific horizon on which he could separate the
Sykes Mountain Formation from the Thermopolis Shale.
In all exposures measured by the writer, there is an abundance of dahllite concretions (calcium phosphate nodules) in a matrix of black shale at the base of the
Thermopolis Shale. They commonly occur as surficial float associated with exposures of the shale. They were developed in place as discussed by McConnell (1935). The writer placed the contact just below the lowest dahllite concretion rich black shale horizon. Furthermore, a dahllite apatite rich bed has been used as the basis for the recognition of a 38
FIGURE 7. Thermopolis Shale (TS) above Sykes Mountain Formation (SMF)
submarine discontinuity in the Middle Devonian of New York by Baird (1978), although he did not find an erosion surface associated with the bed. Therefore, placing the Sykes
Mountain Formation-Thermopolis Shale contact below the dahllite concretion rich horizon is supported by Baird's
(1978) work, since this bed represents a hiatus between the two formations. 39
Regional Correlations
The complex stratigraphy, the proliferation of formational names, extensive facies changes, differing opinions regarding contact relationships between formations and the scarcity of guide fossils as time indicators make regional correlations of Lower Cretaceous sediments difficult. As a result, most of the correlations made by the previous investigators and by the writer are based on the lithologie and genetic similarities or on age equivalence as proven by earlier workers. Many published and unpublished reports were used to establish correlations of the Cloverly-Sykes Mountain-Thermopolis sequence in the
Bighorn Basin with equivalent deposits in adjacent areas.
These reports include Eicher (1960; 1962), Moberly (1956;
1960; 1962), Gries (1962), Haun (1959; 1963), Haun and
Weimer (1960), Haun and Barlow (1962), Weimer (1962), Young
(1960; 1970), and Kvale (1986).
Comparison of the lithologie characteristics of the
Cloverly, Sykes Mountain, and Thermopolis formations with equivalent sediments in central Montana, Black Hills,
Colorado Front Range, Colorado Plateau, and western Wyoming suggests the following correlations.
(1) In central Montana, the Kootenai Formation is the lithologie equivalent of the Cloverly Formation and the 40
Third Cat Creek Sand has the same lithogenetic origin as the
Pryor Conglomerate Member of the Cloverly. The lithogenetic equivalent of the Sykes Mountain Formation is the First Cat
Creek Sand (Moberly, 1956; 1960; 1962).
(2) Comparison of sediments in the Bighorn Basin with those of the Black Hills suggests that the Cloverly
Formation is the lithologie equivalent of the lower part of the Inyan Kara Group (the Fuson-Lakota sequence) which is composed of nonmarine, variegated claystones, siltstones and sandstones interlensed with local conglomerate (Waage,
1958). This sequence underlies the Fall River Formation and is difficult to subdivide and to compare with the members of the Cloverly Formation in the Bighorn Basin (Moberly, 1956).
The variegated claystones and siltstones in the upper Lakota
Formation correlate with the variegated mudstones of the
Little Sheep Mudstone Member of the Cloverly Formation in the Bighorn Basin (Kvale, 1986). According to many previous workers, the transitional marine Fall River Sandstone which is composed of flaggy sandstones, siltstones, and dark shale and which disconformably overlies the Fuson-Lakota sequence is the lithogenetic equivalent of the Sykes Mountain
Formation of the Bighorn Basin. The "Rusty Beds" (Sykes
Mountain Formation) change from west to east; they become thinner on the east side of the Bighorn Mountains and become 41 sandier eastward accross the Powder River Basin, and are composed mainly of sandstone and intercalations of silty shales in the Black Hills (Eicher, 1960; 1962).
Many earlier workers correlated the marine Thermopolis
Shale of the Bighorn Basin with the Skull Creek Shale of the
Black Hills. Eicher (1960; 1962) reported that the Skull
Creek Shale in the Black Hills decreases in thickness from north to south due to the thinning and grading of the upper part of the lower shale and the middle silty shale southward into the upper part of the Fall River Sandstone. As a result, in the southern Black Hills the entire Skull Creek is correlated with the upper shale unit of the Thermopolis in the Bighorn Basin.
(3) The Cloverly Formation in the Bighorn Basin is lithogenetically equivalent of the Lytle Formation of the
Colorado Front Range. The Sykes Mountain Formation is the equivalent of the lowermost member of the South Platte
Formation (Plainview Sandstone Member), and the Thermopolis
Shale is related to the members above the Plainview
Sandstone Member and below the First Sandstone including the
Lower Shale, Kassler Sandstone, and Van Bibber Shale members
(Waage, 1955; 1958; 1961; Young, 1960). The Thermopolis
Shale thins from north to south of the Wind River Basin, and the lower shale and middle silty shale units become obscure farther south. The thinning of the entire section suggests 42 less subsidence in this area than in the adjacent areas during the deposition of the Thermopolis (Richer, 1962).
(4) In the Colorado Plateau, the varicolored floodplain mudstones, conglomeratic sandstones and conglomerates of the
Burro Canyon and Lower Cedar Mountain formations are equivalent to the Cloverly Formation in the Bighorn Basin.
The Lower Naturita Formation (the transitional unit laid down at the time of the earliest Cretaceous marine transgression) is equivalent to the Sykes Mountain Formation
(Young, 1960; 1970). The marine Glencarin Shale Member of the Purgatoire Formation is the equivalent of the
Thermopolis Shale in the Bighorn Basin (Waage, 1955; Young,
1970).
(5) Westward of the Bighorn Basin most of the Gannett
Group including the upper part of the fluvial Ephraim
Conglomerate, the lacustrine Peterson Limestone, the fluvial
Bechler Formation, and the lacustrine Draney Limestone are
Aptian in age as evidenced by microfaunae, and the fluvial
Smoot Formation is Late Aptian-Early Albian in age (Eyer,
1969). The sequence from the upper part of the Ephraim
Conglomerate through the Draney Limestone correlates with the Cloverly and Kootenai formations of Wyoming and Montana
(Eyer, 1969; Furer, 1970). Holm et (1977) believed that the nonmarine Smoot Formation of western Wyoming (Eyer, 43
1969; Durkee, 1980) is the time-stratigraphic equivalent of the Sykes Mountain Formation of the Bighorn Basin. Kvale
(1986, Fig. 25, and Table 4) correlated the upper Ephraim
Conglomerate with the lower Little Sheep Mudstone Member and the Bechler Formation with the upper Little Sheep Mudstone.
Burk (1957) defined the nonmarine interbedded sandstones and finer grained Bear River Formation as the equivalent to the
Lower Cretaceous marine Thermopolis Shale. On the western flank of the Bighorn Basin, the "Rusty Beds" (Sykes Mountain
Formation) are more sandy than on the eastern flank (Burk,
1957; Eicher, 1962). The facies changes similarly from the eastern Bighorn Basin to the east as well as to the west, suggesting that the center of the depositional basin of the
Sykes Mountain Formation was near the eastern Bighorn Basin and the source areas and the environments of higher energy were eastward and westward (Eicher, 1960; 1962). Figure 8 shows the lithogenetic equivalents of the Sykes Mountain
Formation and the adjacent formations.
Summary
The Lower Cretaceous (Albian) Sykes Mountain Formation in the Bighorn Basin consists of the rust-colored, interbedded and interlaminated sandstones and siltstones and dark gray shales that disconformably overlie the fluvial CENTRAL IDAHO- BIGHORN BLACK COLORADO COLORADO AND WYOMING BASIN HILLS FRONT PLATEAU SOUTHWEST BORDER RANGE MONTANA
VAN BIBBER SHALE BEAR THERMOPOLIS SKULL SHALE MEMBER GLBNCAIRN MEMBER RIVER CREEK KASSLER SS. SHALE FORMATION SHALE SHALE MEMBER LOWER SHALE 1
FIRST PNCQ CAT 0 SMOOT SYKES FALL PLAINVIEW LOWER CREEK MOUNTAIN RIVER SANDSTONE NATURITA SAND FORMATION FORMATION formation MEMBER FORMATION
STONE 1 GROUP
draney GREYBULL UPPER lmst. g SB. LAKOTA CEDAR MOUNTAIN g L. MIMES § FORMATION BECHLER FUSON u. LITTLE g S FORMATION SHEEP 5 MUDSTON E < PETERSON DK CHERT E4 BUCKHORN GANNETT KOOTENAI FORMATION FORMATION LOWER ^ 1 LMST. > CONGL. § LAKOTA CONGLOMERATE UPPER S L LITTLE 3 EPHRAIM o SHEEP MUDST.
FIGURE 8. Lithoqenetic equivalents of the Sykes Mountain Formation and the adjacent formations 45 nonmarine Cloverly Formation and grade upward into and interfinger with the marine Thermopolis Shale. The thickness of the Sykes Mountain Formation is variable as a result of its interfingering with the overlying Thermopolis
Shale. The Sykes Mountain Formation includes the Rusty
Beds, Dakota Silt, and Dakota Sandstones of some previous workers. The entire formation is often stained with iron oxide and contains abundant ironstone concretions and laminae. Sandstones are dominant in the lower part of the formation, whereas the shales increase upward, forming an overall fining upward sequence. The formation exhibits a suite of primary structures including different kinds of stratification, ripples, and biogenic structures.
Generally, the Sykes Mountain Formation is rarely fossiliferous, except for a few poorly preserved pelecypod fossils that occur in the sandstones near the base of the formation in some locations. The Sykes Mountain Formation represents the transitional marine environment of the Early
Cretaceous transgressive sea. It is correlated with the
First Cat Creek Sand of central Montana, the Fall River
Formation of the Black Hills, the Plainview Sandstone Member of the Colorado Front Range, the Lower Naturita Formation of the Colorado Plateau, and the nonmarine Smoot Formation of western Wyoming. 46
CHAPTER II. FACIES AND DEPOSITZONAL ENVIRONMENTS
Introduction
The Sykes Mountain Formation of the Bighorn Basin is
Early Cretaceous (Albian) in age. It is mainly composed of a sequence of the clastic sediments which disconformably overlie the nonmarine sediments of the Cloverly Formation and conformably underlie the marine black shales of the
Thermopolis Formation. It grades upward into the
Thermopolis Shale and interfingers with it, causing a variation in the thickness of both formations. It is generally poorly exposed, but there are some excellent outcrops within the Bighorn Basin along Sheep Mountain
Anticline, Rose Dome, Little Sheep Mountain Anticline, Sykes
Mountain, Red Pryor Anticline, in the Beaver Creek area, in the vicinity of Hyattville, north of Thermopolis, and west of Cody. Exposures were measured and described at each of these areas during the summers of 1984 and 1985 (Figure 1 and Table 1).
The Sykes Mountain Formation records the initial encroachment of the Early Cretaceous sea in the Western
Interior. As far as the writer is aware, there is no detailed work regarding a facies analysis of the Sykes
Mountain Formation in the Bighorn Basin. However, several 47
authors have studied the stratigraphy, paleontology and
petrology of the Sykes Mountain Formation, considering it
either as a separate formation or as the upper part of the
Cloverly Formation or the lower unit of the Thermopolis
Shale according to the different stratigraphie nomenclatures
used by each author.
The objectives of this chapter of this report is to
analyze the different facies within the Sykes Mountain
Formation and to attempt to relate the various facies
associations to the depositional environments they record.
The writer tries to establish these objectves by answering
the following questions:
(1) What are the different facies present within the
Sykes Mountain Formation?
(2) What are their characteristics?
(3) How are they related to each other and what facies
associations can be identified?
(4) What are the possible environments of deposition
recorded by the facies associations?
To answer these questions, the writer measured several
stratigraphie sections within the Bighorn Basin and
described the different lithologies, including the various
primary structures present, within the Sykes Mountain
Formation. 48
Description of Faciès
Lithology and primary structures are used to interpret distinct depositional environments and processes of deposition. Code names are erected for the lithofacies of the Sykes Mountain Formation following Miall's (1977; 1978) lithofacies code system. Each code name consists of a capital letter referring to the dominant grain size followed by lower case letters designating the main primary structures.
The sediments of the Sykes Mountain Formation are divided into two lithofacies associations: (1) the Quartz
Arenite Facies Association (QA) and (2) the Heterolithic
Facies Association (SM). The QA Facies Association is recognized within the sandstone bodies and is divided into seven facies on the basis of differences in sedimentary structures. These seven facies, named for their common sedimentary structures are; (1) Flat-Bedded and Parallel-
Laminated Sandstone Facies (Sfl), (2) Planar Cross- stratified Sandstone Facies (Sp), (3) Trough Cross- stratified Sandstone Facies (St), (4) Hummocky Cross- stratified Sandstone Facies (She), (5) Convolute-Bedded
Sandstone Facies (Scb), (6) Asymmetrical-Rippled Sandstone
Facies (Sar), and (7) Symmetrical-Rippled Sandstone Facies
(Ssr). The SM Facies Association is represented by the 49 sediments which occur between the sandstone bodies of the QA
Facies Association. The main characteristics of these facies are represented in Table 2, and further information is given below.
Quartz Arenite Facies Association (QA)
The QA Facies Association consists of yellow and rust- colored, fine- to very fine-grained, quartz arenite.
Sandstones of the QA Facies Association which occur near the top are finer-grained than those which occur near the base of the Sykes Mountain Formation. The thickness of the QA
Facies Association varies from one to two meters in the lower parts of the formation to a few centimeters in the upper parts. In addition, the thickness of the QA Facies
Association varies from one location to another. The QA
Facies Association contains ironstone laminae and concretions, and pyrite nodules occur in some places. This facies association possesses a sharp, flat, erosive base, and a sheet-like geometry and is laterally persistent and traceable for long distances. It grades upward to the SM
Facies Association. In the Beaver Creek area, the QA Facies
Association is only 20 to 30 cm thick, and the rest of the
Sykes Mountain Formation is represented by the SM Facies
Association. Sandstones of the QA Facies Association are TABLE 2. Lithofacies characteristics and their probable depositional environments
Facias Facies Lithology Sedimentary Interpretation Assoc. structures
QA Sfl Sandstone, fine Flat bedding and parallel Low to middle to very fine, lamination, burrows and sand flats Fe-oxiae stained trails Sp Sandstone, fine Large-scale planar Subtidal and to very fine cross-stratiiication, intertidal herringbone cross- environments stratification, reactiva tion surfaces, mud drapes, burrows and trails St Same as Sp Large-scale trough Subtidal to cross-stratification, intertidal burrows She Same as Sp Hummocky cross- Shallow and coquinas stratification, marine shelf in places burrows Scb Sandstone, fine Convolute bedding Tidal flat Sar Sandstone, very Asymmetrical ripples, Intertidal f ine dunes, burrows, trails Ssr Same as Sar Symmetrical ripples, Intertidal flat-topped ripples, flats burrows, and trails SM Sandstone, Flaser-, wavy-, and Shallow sub siltstone, lenticular bedding, tidal end and mudstone burrows and trails intertidal 51 well indurated and cemented by silica, iron oxides, or carbonates. They contain poorly preserved pelecypod valves and plant fragments in some places. The QA Facies
Association exhibits a variety of primary structures. It is divided into seven lithofacies based on their respective sedimentary structures.
Flat-Bedded and Parallel-Laminated Sandstone Facies (Sf1)
The Sf1 Facies consists of well indurated, fine- to very fine-grained quartz arenite. It is cemented by silica and iron oxides. Sandstones of the Sf1 Facies are noncalcareous near the base of the formation and calcareous in the upper parts. The facies overlies either the variegated mudstone of the Himes Member of the Cloverly
Formation or the interbedded sandstones, siltstones, and mudstones of the SM Facies Association. It also may overlie
Facies Sp, St, Scb, Sar, or Ssr and underlie Facies Sp, St,
She, Sar, or Ssr with sharp contacts. Sandstones of this facies have a sharp erosive base and are laterally persistent. However, some of the sandstone bodies, especially near the base of the section are lenticular in shape. The Sf1 Facies varies in thickness from about 50 cm in the lower parts of the formation to a few centimeters near the top. However, thickness of the sandstone bodies near the base of the formation varies from one location to 52 another. In addition to the thickness change, grain size varies from the base to top. In every section, sandstones of the Sfl Facies present near the base are thicker and consist of coarser grains than those in the upper parts of the Sykes Mountain Formation. Ironstone beds and pyrite nodules occur within the Sfl Facies in some locations.
The important primary structures in the Sfl Facies include flat bedding and parallel lamination (Figure 9) with parting lineation. The Sfl Facies contains mud drapes between its laminae. Burrows and trails occur in and on the sandstones of this facies. In the Beaver Creek area, the QA
Facies Association is only represented by 20 to 30 cm of brown and dark gray, very fine-grained sandstone and siltstone of the Sfl Facies. At Section RD. 12 at Rose Dome
(Table 1), and near the bottom of the formation, the Sfl
Facies is dark brown, marked by an ironstone bed at its base, contains shell fragments (coquina) at its top, changes laterally to a quartz arenite and then pinches out over a very short distance. The Sfl Facies contains poorly preserved pelecypod valves, shell fragments and silicified wood and plant fragments at many places of the area of study. The pelecypod valves are usually convex up. 53
FIGURE 9. Flat bedding (FB) in the Sfl Facies underlying hununocky cross-stratification (HCS) in the She Facies
Planar Cross-stratified Sandstone Facies (Sp)
This facies occurs in several locations in the lower part of the Sykes Mountain Formation. The Sp Facies ranges in thickness from a few centimeters to 50 cm. The beds and laminae are laterally persistent over tens of meters. The
Sp Facies is composed of a fine- to very fine-grained, iron- oxide stained quartz arenite. It may or may not be calcareous. The most important primary structures are large-scale planar cross-stratification (Figure 10) and large-scale herringbone cross-stratification (Figure 11) 54 with sharp set boundaries. The planar cross-stratification has an average set thickness of 25 cm and angles of inclination of 19 degrees. The herringbone cross- stratification has an average set thickness of 20 cm and inclination angles of 15 degrees. Some of the cross- stratification foresets are truncated by reactivation surfaces and are draped by mud. The sandstones of this facies have sharp, commonly flat, and sometimes undulatory bases and tops. Burrows and trails are present. Sandstones of the Sp Facies may overlie and underlie Facies Sfl, Sar,
Ssr, or the SM Facies Association. Figure 12 shows the asymmetrical rippled sandstone of the Sar Facies overlying the cross-stratified sandstone of the Sp Facies. The ripple marks above and the planar cross-stratification below indicate opposite directions of paleocurrent. Furthermore, on the upper surface of the sandstone body present at the base of the section measured at Red Pryor Anticline (Section
RPA. 17, Table 1), sandwaves with straight crests, and a wavelength of 1.80 m and amplitude of 30 cm have been recorded (Figure 13).
Trough Cross-stratified Sandstone Facies (St)
Sandstones of the St Facies have the same lithologie characteristics of those of the Sp Facies. The most important primary structure in this facies is the large- 55
FIGURE 10. Large-scale planar cross-stratification (PCS) in the Sp Facies
scale trough cross-stratification (Figure 14) which is believed to be produced by the migration of dunes. The trough cross-stratification usually has a set thickness of more than 5 cm and a width of up to 2 meters. This facies has sharp boundaries, and its base is usually marked by a scoured surface. Biogenic structures including burrows and trails are common in the St Facies. Sandstones belonging to the St Facies occur near the base of the Sykes Mountain
Formation only in a few localities scattered throughout the
Bighorn Basin (at Section RD. 12, Section LSMA. 8, Section
SM. 14, and Section C. 24, Figure 1 and Table 1). 56
m
FIGURE 11. Herringbone cross-stratification in the Sp Faciès
Sandstones of the St Faciès may overlie the Sfl, Sp Facies, or the SM Facies Association and underlie the Sar, Ssr, Sp
Facies, or the SM Facies Association.
Hununocky Cross-stratified Sandstone Facies (She)
The average thickness of the hununocky cross-stratified sandstone of the She Facies is 25 cm. Lamination within the sandstone of this facies is generally well defined and the laminae are less than 1 cm thick. It occurs in several localities throughout the area of study, especially in the 57
FIGURE 12. Planar cross-stratified sandstone of the Sp Facies below the rippled sandstone of the Sar Facies
lower part of the formation. It overlies and underlies the flat-bedded sandstones of the Sfl Facies and also may underlie the rippled sandstones of Facies Sar or Ssr with either gradational or sharp erosional contacts. The hummocky cross-stratified sandstone consists of fine- to very fine-grained quartz arenite in most of the places at which it occurs. However, in a few locations, coquinas consisting of pelecypod valves and silicified wood fragments occur. An example of the hummocky cross-stratification is shown in Figure 9. 58
FIGURE 13. Sandwaves in the Sp Facies at Red Pryor Anticline
Convolute-Bedded Sandstone Facies (Scb)
The Scb Facies is composed of fine-grained quartz arenite and is characterized by the presence of convolute bedding in the form of anticlines and synclines. The convolution increases in intensity upward within this facies and is sharply truncated at the top. The thickness of the
Scb Facies is usually about 30 cm. The Scb Facies has been observed in a few localities scattered throughout the area of study (at Cody, Thermopolis, and Red Pryor Anticline).
It occurs only in the lower part of the Sykes Mountain 59
FIGURE 14. Large-scale trough cross-stratification in the St Facies
Formation. The Scb Facies overlies and underlies the parallel-laminated sandstones of the Sfl Facies with sharp and erosional contacts. Figure 15 shows an example of the convolute bedding in the Scb Facies.
Asymmetrical-Rippled Sandstone Facies (Sar)
The Sar Facies is also composed of iron-oxide stained, very fine-grained quartz arenite. It occurs in most measured sections. The main primary structure is the asymmetrical ripple marks with common straight crests 60
FIGURE 15. Convolute bedding in the sandstones of the Scb Facies
(Figure 12). The ripple marks are usually mud draped. Wave lengths of the asymmetrical ripples are variable, with an average value of 15.1 cm, and the heights are less than 5 cm. The Sar Facies is generally in the range of a few centimeters thick. Biogenic structures including vertical and horizontal burrows, trails, and other unidentified structures are abundant on and near the surface of the sandstones of the Sar Facies (Figure 16). These biogenic structures may have destroyed other primary structures partly or completely in some places. These burrows, trails, 61 and other biogenic structures may be due to the action of the pelecypods and/or any other unidentified organisms at the time of or shortly after the deposition of these sediments. The Sar Facies usually overlies Facies Sfl, Sp,
St, or She and underlies Facies Sfl, Sp, St, or the SM
Facies Association.
FIGURE 16. Biogenic structures including burrows and trails on Facies Sar and Ssr
Symmetrical-Rippled Sandstone Facies (Ssr)
The Ssr Facies has the same lithology as the Sar Facies and occurs in most Sykes Mountain Formation outcrops. This facies is characterized by the presence of symmetrical 62 ripples with straight crests and wave lengths with an average value of 12.6 cm (Figure 17). The symmetrical ripple marks sometimes have flat or rounded crests. Burrows and trails are also common in this facies. The Ssr Facies also may overly Facies Sfl, Sp, St, or She and underlie
Facies Sfl, Sp, St, or the SM Facies Association.
FIGURE 17. Straight-crested symmetrical ripples in the Ssr Facies 63
Heterolithic Faciès Association (SM)
The SM Facies Association consists of mixed lithologies of rusty brown, interbedded, very fine-grained sandstones, siltstones, and silty shales. The sandstones and siltstones exhibit different colors like those of the different sandstone facies, and the shales are either gray or black.
The silty shales are very thinly laminated or fissile with increasing fissility and dark color upsection. Ironstone laminae and concretions are common. Sandstones and siltstones are in the form of very thin, discontinuous laminae or in the form of lenses within the mudstone. Cone- in-cone concretions occur at the base of this facies association in some areas. The sandstones and siltstones may be ferruginous or calcareous. Black and dark gray carbonaceous plant (leaf and wood) fragments commonly occur within this facies association. Within the entire formation, sandstones and siltstones are dominant in the lower parts, with shales increasing in the upper parts. The
Heterolithic Facies Association represents fining upward sequences, in which shale increases in amount upward. This facies association occurs in association with one facies or more of the QA Facies Association in a fining upward sequence which is repeated several times within the formation. The SM Facies Association, like Sfl Facies, is 64 finer in the Beaver Creek area and becomes coarser to the west of this area.
The primary structures present in these mixed lithologies of sandstone and mudstone are similar to those described by Reineck and Wunderlich (1968) and include lenticular bedding, wavy bedding, flaser bedding, and thinly and thickly alternating sandstone and mudstone layers. The sand layers are either current or oscillation ripple structures. Figure 18 shows an example of wavy and lenticular bedding present in this facies association.
Biogenic structures are common and are represented by burrows and trails. The SM Facies Association usually underlies the sandstones of Facies Sfl, Sp, St, or Scb and overlies Facies Sfl, Sp, or St. The uppermost sequence of the SM Facies Association conformably underlies the marine black shales of the Thermopolis Formation. It grades into and interfingers with the Thermopolis Shale causing variation in the thickness of both Sykes Mountain and
Thermopolis formations. The contact of the uppermost sequence of this facies association with the overlying
Thermopolis Shale is characterized by the presence of a marker horizon rich in dahllite concretions which occurs at the base of the Thermopolis Shale. 65
FIGURE 18. Lenticular bedding (LB) and wavy bedding (WB) i the SM Facies Association
Facies Sequence
The Sykes Mountain Formation facies are grouped into sequences. Each sequence consists of the QA Facies
Association overlain by the SM Facies Association. The QA
Facies Association has a sharp contact with the underlying deposits of the SM Facies Association. The QA Facies
Association grades upward into the SM Facies Association, forming a fining upward sequence. There are at least four fining upward sequences composed of alternating QA and SM 66 facies associations. Figure 19 shows examples of the vertical sequences in the Sykes Mountain Formation, and
Figure 20 shows the alternation between the QA and SM facies associations.
Interpretation of the Depositional Environments
Lithology and sedimentary structures present in the different facies as well as the facies sequences of the
Sykes Mountain Formation are used to interpret the environments of deposition of these sediments. Flat beds form in medium- and fine-grained sand under the conditions of upper flow regime (high flow velocity and shallow water depth) where ripples and dunes are washed out (Collinson and
Thompson, 1982). However, Harms et (1975, 1982) showed that plane beds may also develop under lower flow regime conditions in sand coarser than 0.6 mm and under upper flow regime conditions in all sand sizes. They form as a result of unidirectional currents, oscillatory currents and also by combined flow (Allen, 1982; Harms et , 1982). Bartsch-
Winkler and Ovenshine (1984, p. 1235) stated that: "Planar bedding is characteristic of the high-flow regime during the initial flood stage and midstage flood and ebb flow
(sediments of the low and middle flats)." The change of a succession from ripples to plane beds with erosional contact 67
SM SM ] 7S /N- /\ /\ /% I V/l W VJ
SM SM
SM SM I SM
/A
SM
SM
SM
Flat Bedded Ss. of Facias Sfl Cross Stratified Ss. of Fades SPt I f Heteroiithic Lithologie* of Herringiioiie Cross-St ratified I I Fades Assoc. SM Sa. of Faciei Sh n Reactivation Surface# and Hutnmcclcy cross.stratified Mud Drape* Ss. of Facias Sk Burrow* Convoluto Bedded Ss. m ïmm of Facia* Sc Rippled Ss. of Facias 1^ A Ssr or Sar
FIGURE 19. Examples of vertical sequences in Sykes Mountain Formation 68
FIGURE 20. Alternation between the QA and SM facies associations
suggests an abrupt increase in flow strength from that of the lower flow regime to that of the upper flow regime, but the transition from plane beds to ripples indicates a gradual change from upper to lower flow regime (Ly, 1982, p.
201). Banks (1973, p. 217) reported that the occurrence of cross-lamination overlying parallel lamination indicated an upward decrease in current strength. Moreover, Reineck
(1967b, p. 193) reported that this type of sedimentation which results in laminated sand is very common in beach sediments. Osgood (1970) and Simpson (1975) suggested that the single vertical burrows are the results of movement of 69 organisms escaping rapid sedimentation or searching for food within the sediments. In addition to the trace fossils, lag concentrations of mollusk shells and coquinas in middle and high tidal flat environments have been reported by Van
Straaten {1954b), Evans (1965), and Klein (1963; 1967).
Pelecypod valves lie dominantly with convex sides up due to the action of waves and currents in an intertidal setting
(Van Straaten, 1952).
Large-scale tabular and trough cross-stratification structures are produced from the migration of large scale two-dimensional and three-dimensional ripples, respectively
(Harms e_t al., 1982). Both tabular and trough cross- stratification indicate unidirectional flow. Cross- stratification, with sharp upper and lower set boundaries, characterizes intertidal sand bodies formed by tidal current flow (Klein, 1970a; Swett et , 1971). Moreover, a tidal origin is indicated by the occurrence of reactivation surfaces and mud drapes (Collinson, 1970; Klein, 1970a;
1970b). The common occurrence of burrowing structures supports a tidal origin (Rhoads, 1967). Further evidence of tidal origin comes from the bimodal-bipolar distribution of cross-stratification orientation where ebb and flood operate
(Swett et , 1971). King and Chafetz (1983, p. 269) reported low-angle cross-stratification sandstone representing a major accumulation of tidal flat. 70
Herringbone cross-stratification is a bidirectional current-
formed sedimentary structure in which sets of opposite
dipping cross-strata reflect the to and fro movement of ebb
and flood tidal flows (Johnson, 1978; Selley, 1982). It is
common in tide-dominated environments of the German North
Sea (Reineck, 1963) and is an indicator of tidal processes
(Swett et , 1971). Herringbone cross-stratification is
one of the most diagnostic features of tidal currents
(Johnson, 1978) and very common in shallow marine
environments (Selley, 1982). Furthermore, Reineck (1975)
reported that herringbone structures and bipolar cross-
stratification characterize the subtidal zone.
Reactivation surfaces are erosional surfaces that form
within a cross-stratified set as consequences of
unsteadiness in flow velocities (Collinson and Thompson,
1982; Mowbray and Visser, 1984). They can be produced in
either unidirectional or bidirectional flow systems (Mowbray
and Visser, 1984). Klein (1970a) interpreted the
reactivation surfaces as a result of morphological
modification of bedform by a weaker reversing tidal current.
Reactivation surfaces are a common and important factor in
the recognition of tidal deposits (Klein, 1970a; Swett et
al.. 1971; Mowbray and Visser, 1984), although they also may
be produced in rivers as a result of discharge fluctuations 71
(Collinson, 1970), and in aeolian deposits (Brookfield,
1977).
Mowbray and Visser (1984) reported that mud drapes are
depositional features which can be used in addition to
reactivation surfaces as a way to recognize the unsteadiness
in flow velocities of tidal processes. According to many
authors (e.g., Reineck and Wunderlich, 1968; Klein, 1970b;
Mowbray and Visser, 1984), large-scale sandy foresets can be
formed during the strong flow of the ebb current while mud stays in suspension. They added that the mud drapes are
formed during the slack water period of the tidal cycle when
the flow velocity is too low to allow mud to be deposited
from suspension. On the contrary, McCave (1970) reported
that the short period of slack water is not long enough for
these mud drapes to be deposited from suspension because the
concentration of suspended sediment and the settling
velocity are low. Therefore, he attributes their formation
to the result of a combination of abnormally high
concentrations of suspended fine-grained material, low wave
intensity, and low current velocity over a longer period.
Mowbray and Visser (1984) assume that biological factors are
very important in mud drape formation in increasing settling
velocities by flocculation and pellitization of fine-grained
material. Mud drapes are common in tidal deposits (Reineck, 72
1963; Raaf and Boersma, 1971), and are commonly associated with tidal sand waves (Terwindt, 1971; Bouma et , 1980).
Harms et (1975) introduced the term hummocky cross- stratification or hummocky bedding. It is thought to be formed by deposition over a hummock-and-swale surface during the oscillatory motion of storm waves (Harms et , 1975;
1982). "Much of the hummocky cross-stratification results from the filling of scours within or at the base of the hummocky cross-stratified sandstone" (Hunter and Clifton,
1982). It is most common in sediments of shallow marine shelf origin and is believed to result from the action of storm waves on the bottom of the shelf at depths below normal, fair weather wave reworking (Collinson and Thompson,
1982). Harms et (1975) attributed hummocky cross- stratification to a flow of higher velocity than that needed for ripple formation. Its presence is used to identify the inner shelf and lower shoreface zones in shallow marine sequences (Walker, 1984).
Convolute bedding consists of folding of lamination with sharp anticlines and gentle synclines. It is a soft sediment deformational structure that may occur in deposits of many different sedimentary environments. It may be produced in several different ways, but the liquefaction of the sediment shortly after deposition is the most important 73 factor in the genesis of convolute bedding (Reineck and
Singh, 1980). It could also be initiated by cyclic wetting and drying in the intertidal zone (Doe and Dott, 1980).
"The main use of convolute lamination is as evidence of rapid deposition" (Collinson and Thompson, 1982, p. 145).
Convolute bedding and lamination are common "just below the sediment surface in present-day river floodplains and tidal flats" (Collinson and Thompson, 1982).
Ripples form commonly in coarse silt and fine- to medium-grained sand subjected to gentle traction currents in the lower part of the lower flow regime (Collinson and
Thompson, 1982; Selley, 1982). Ripples are related to certain flow conditions which may occur in several different environments. Clifton (1983) reported that sand waves and megaripples are formed by tidal currents, whereas ripples are developed in estuarine sand by waves on intertidal flats or by tidal currents in or adjacent to tidal channels.
Asymmetrical ripples indicate deposition by unidirectional currents. Symmetrical and straight crested ripples indicate oscillatory conditions during deposition (Collinson and
Thompson, 1982; Hunter and Clifton, 1982). Banks (1973, p.
217) reported that the occurrence of symmetrical ripples indicates deposition in the zone of effective wave action.
Flat-topped ripples with eroded crests are formed during 74 modification by falling water levels and a change in direction of flow (Van Straaten, 1954a; Tanner, 1958; Klein,
1971). Migration of ripples produces cross-lamination
(Collinson and Thompson, 1982). Dunes and sandwaves are similar to current ripples in many features but on a larger scale. They are common in deposits of tidal channels and rivers (Collinson and Thompson, 1982). Currents responsible for sandwaves and dunes are stronger than those that generate ripples. The occurrence of mud drapes over current ripples may suggest an evidence of tidal origin.
Trace fossils are important in interpreting environments of deposition. They reflect the animal behavior and responses toward environmental changes (Rhoads,
1975). The presence or absence and the degree of bioturbation are related to the rate of deposition (Reineck,
1967b). Bioturbation is abundant in fine-grained sediments and indicates rapid rate of sedimentation in sediments deposited under tidal conditions (Reineck and Singh, 1980; van Straaten, 1954b). Deep vertical burrowing is a characteristic of near-shore especially intertidal environments, and shallower horizontal burrowing to depths of 10 cm or more is developed in sediments farther offshore
(Rhoads, 1967). The sedimentation rates can be indicated by burrowing intensity (Moore and Scruton, 1957). Slow, 75
continuous sedimentation may be indicated by intense
burrowing and complete bioturbation of the sediment that
destroys stratification, whereas rapid, continuous
sedimentation is reflected by slight or no burrowing (Ekdale
et al., 1984). However, "highly burrowed sandstones may not
have been deposited in low energy environments with low
rates of sedimentation" (Miller, 1984, p. 209). Miller
(1984) added that the biogenic structures which occur in
quartz arenites may have resulted from deposit-feeding or
other active animals. Many workers (e.g., Rhoads, 1967;
Whitlatch, 1980) reported the abundance of the deposit-
feeders in intertidal sandflats.
The environment of deposition of the Heterolithic
Facies Association was one of fluctuating current strengths,
as indicated by the alternation of sandstones, siltstones,
and mudstones. The structures which characterize the mixed
bedding or mixed lithologies found in tidal environments
include flaser, wavy, and lenticular bedding (Van Straaten
and Kuenen, 1958; Reineck, 1960a; 1967b; Reineck and
Wunderlich, 1968; Reineck and Singh, 1980). These
structures may be formed in other environments as well.
Flaser bedding "implies sand and mud, as well as current
activity and pauses in this activity" (Reineck, 1967b, p.
195). Sand is transported and deposited in the form of 76
ripples while the mud is held in suspension during current
activity times. The mud which is held in suspension is deposited and covers the ripple troughs or completely covers the ripples during slack water. At the beginning of the
following current cycle the mud is either eroded from the crests of the sand ripples and remains only in the troughs or is covered by newly deposited sand (Reineck, 1967b).
Flaser bedding is either simple or wavy. In simple flaser bedding the mud is deposited and preserved in the ripple troughs without contacts between the mud flasers, while in wavy flaser bedding the mud flasers also cover the ripple crests with discontinuous layers (Reineck and Wunderlich,
1968). In wavy bedding, the mud layers fill the ripple troughs and cover the ripple crests. Pasierbiewicz (1982) proposed a new concept of flaser and wavy bedding genesis.
According to Pasierbiewicz's (1982) hypothesis, the mud of the tidal environment mixed bedding does not only develop when the current activity pauses, but it can also be formed during the tidal current activity. Lenticular bedding consists of sand lenses (isolated oscillation or current ripples) of irregular form embedded in a mud groundmass
(Reineck, 1967b). The sand lenses are either single or connected (Reineck and Wunderlich, 1968). The alternating sand-mud layers consist of slightly undulating, parallel- 77
bedded sand and mud layers (Terwindt and Breusers, 1972).
Such structures are mainly observed in the shallow sub-tidal
and intertidal environments (van Straaten, 1954a; Reineck
and Wunderlich, 1968). They indicate constant fluctuating,
relatively low-energy conditions, with short periods of alternating sand and silt transport by tidal currents and waves with mud deposition occurring from suspension (Reineck and Wunderlich, 1968).
The above discussion of the lithofacies and their characteristics and the processes responsible for the
formation of sedimentary structures of each respective
facies, as well as the criteria used by different workers, suggest that the flat-bedded and parallel-laminated quartz arenites of the Sfl Facies with convex upward pelecypod valves and vertical burrows represent sediments of the low and middle flats formed under the conditions of upper flow
regime. The Sp Facies is interpreted as being intertidal
and subtidal deposits on the basis of its primary structures
including planar cross-stratification with sharp set
boundaries, herringbone cross-stratification, reactivation
surfaces, mud drapes, sand waves, common burrowing structures, and bimodal-bipolar distribution of the cross- stratification orientation. Sediments of the St Facies were
also interpreted as being deposited in subtidal and 78 intertidal environments due to the presence of trough cross- stratification with sharp set boundaries and the bimodal- bipolar distribution of the cross-stratification. Presence of hummocky cross-stratification in the She Facies reflects the action of storm waves on shallow marine shelf sediments.
Sediments of the Scb Facies are interpreted as being tidal flats deposited under rapid sedimentation rate conditions, as indicated by the presence of the convolute bedding.
Presence of asymmetrical ripples, mud drapes over the ripples and vertical burrows in facies Sar suggest its tidal origin produced by tidal currents. Facies Ssr is interpreted as intertidal flats and tide-dominated deposits that were produced by tidal waves, as indicated by symmetrical ripples, flat-topped and rounded crest ripples, burrows and trails. Primary structures including flaser, wavy, and lenticular bedding which occur in the SM Facies
Association are mainly observed in shallow subtidal and intertidal environments.
The fining upward sequence which consists of the Quartz
Arenite Facies Association (QA) and the Heterolithic Facies
Association (SM) with the vertical change in sedimentary structures and general upward increase in burrows number suggests an upward decrease in energy. Comparison of the fining-upward sequence and its sedimentary structures with 79 the sequences reported from modern and ancient documented tidal flat and deposits of tide-dominated environments (Raaf and Boersma, 1971; Klein, 1963; 1967; 1970a; 1970b; 1971;
1975; MacKenzie, 1968; Reineck, 1963; 1967b; 1975; Van
Straaten, 1954a; 1954b; 1961; and many others), suggests that the deposits of the Sykes Mountain Formation were deposited in a tidal environment. Tidal processes are indicated by the herringbone cross-stratification, reactivation surfaces, mud drapes, bimodal-bipolar distribution of foreset inclination azimuths. The sequence of parallel lamination with erosional base, hummocky cross- stratification, ripple marks and burrows and trails at the top, and the presence of shell coquinas indicate deposition under the influence of storm waves on an inner shelf. This sequence resembles the sequences described by Craft and
Bridge (1987), Brenchley (1985), Walker et (1983), Harms et al. (1982), Dott and Bourgeois (1982), and many others; with some variations from the ideal hummocky cross- stratification model, such as the common absence of the basal lag unit and the upper parallel lamination unit. The sequence of the convoluted beds separated by flat-bedded and parallel-laminated units resembles those sequences described by several workers (e.g., Reineck and Singh, 1980; Bartsch-
Winkler and Schmoll, 1984), indicating high sedimentation 80
rates and syndepositional or post-depositional deformation
in the high-energy tidal regime.
Comparison of Sedimentary Structures in Sykes Mountain
Formation to Other Tidal Sediments
The Sykes Mountain Formation exhibits several sedimentary features that commonly occur in many Recent and
ancient documented examples of clastic tidal and tide- dominated deposits. Table 3 is a summary of the sedimentary
features in the Sykes Mountain Formation and in some other
Recent and ancient tidal sediments. Comparison of the
assemblage of the sedimentary structures in the Sykes
Mountain Formation with those reported from Recent and
ancient tidal deposits provides a convincing evidence of the tidal origin of these sediments.
Summary
Two facies associations are recognized in the Sykes
Mountain Formation; the Quartz Arenite Facies Association
(QA) and the Heterolithic Facies Association (SM). The QA
Facies Association is divided into seven facies, named for
their common sedimentary structures. The different
lithologies, sedimentary structures of different facies and
their relationships as well as the criteria used by several 81
TABLE 3. Sedimentary features in the Sykes Mountain Formation compared to other tidal sediments
Sykes Mountain Other tidal sediments Formation
Flat bedding and Characteristic of sediments of parallel the low and middle flats formed lamination under high flow regime conditions (Bartsch-Winkler and Ovenshine, 1984)
Large-scale cross- Characteristic of intertidal stratification with sand bodies formed by tidal sharp set boundaries current flow (Klein, 1970a; 1970b; Swett et al., 1971)
Herringbone cross- Common in tide-dominated stratification environments (Reineck, 1963), and is an indicator of tidal processes (Swett et al., 1971; Johnson, 1978; Selley, 1982)
Reactivation surfaces Common and important factor in the recognition of tidal deposits (Klein, 1970a; 1970b; Swett et al., 1971; Mowbray and Visser, 1984)
Mud drapes Common in tidal deposits (Reineck, 1963; Raaf and Boersma, 1971)
Convolute bedding Common in tidal flats (Collinson and Thompson, 1982)
Ripple marks of all Produced by tidal currents types in or adjacent to tidal channels or by tidal waves on intertidal flats (Clifton, 1983)
Flaser, wavy, and Mainly observed in the lenticular bedding shallow subtidal and intertidal environments (Van Straaten, 1954a; Reineck and Wunderlich, 1968) 82
TABLE 3. (Continued)
Sykes Mountain Other tidal sediments Formation
Biogenic Indicate rapid rate of structures sedimentation in sed including iments deposited under burrows and tidal conditions trails (Reineck and Singh, 1980; Van Straaten, 1954b; Rhoads, 1967)
Bimodal-bipolar Common in tidal distribution of environments where ebb cross-stratification and flood tidal currents dip azimuth operate (Swett et al., 1971)
Hummocky cross- Common in shallow marine stratification shelf deposits (Collinson and Thompson, 1982; Walker, 1984)
Convex-up shells Dominant in intertidal and shell fragments deposits (Clifton, 1983; Van Straaten, 1952)
workers have been used to interpret the environments of deposition of the Sykes Mountain Formation sediments.
The Sfl Facies consists of thinly bedded and laminated fine- to very fine-grained, rust-colored, iron-oxide stained quartz arenite, with abundant ironstone laminae and concretions. Poorly preserved pelecypod fossils with convex 83 sides up occur in some locations and vertical burrows are dominant. This facies is interpreted as low to middle sand flats. Facies Sp and St with the same lithology as facies
Sfl are interpreted as intertidal and subtidal deposits due to the occurrence of planar and trough cross-stratification with sharp boundaries, herringbone cross-stratification, reactivation surfaces, mud drapes, burrowing structures, and bimodal-bipolar distribution of cross-stratification set azimuths. The She Facies consists of shallow marine shelf sediments formed under the conditions of storm waves. The
Scb Facies with convolute bedding and lamination is interpreted as a tidal flat. Facies Sar with asymmetrical ripple marks, mud drapes over the ripples, and vertical burrows represents a tide-dominated environment. Facies Ssr with symmetrical ripples, flat-topped and rounded-crest ripples, burrows and trails represents intertidal flats and is tide-dominated.
The SM Facies Association consists of interbedded sandstone, siltstone, and silty shale with ironstone laminae and concretions. Plant fragments and carbonaceous materials are common. The dominant primary structures include flaser, wavy, and lenticular bedding, and biogenic structures including burrows and trails. The SM Facies Association is interpreted as shallow subtidal and intertidal. The QA and
SM facies associations are arranged into sequences. Facies 84
Association QA has a sharp contact with the underlying and gradational contact with the overlying sequence of Facies
Association SM forming a fining upward sequence. Comparison of this fining upward sequence with those reported from
Recent and ancient tidal flat and tide-dominated deposits suggest tidal depositional environment for the Sykes
Mountain Formation sediments. 85
CHAPTER III. PALEOCURRENT ANALYSIS
Introduction
Sedimentary structures are widely used in recognizing certain environments of deposition and in reconstructing paleogeography. They are useful, to some extent, in the determination of the regional paleoslopes (Potter and Olson,
1954) and source areas (Potter and Siever, 1956; Singler and
Picard, 1979), and in reconstruction of the trends of the ancient shorelines (Tanner, 1955).
The general stratigraphy of the Sykes Mountain
Formation has been described by several authors (e.g.,
Eicher, 1960; 1962; Moberly, 1956; 1960; 1962; and the first chapter of this report). Primary structures and facies analysis have been discussed in Chapter II. Several workers
(e.g., Eicher, 1960; 1962; Moberly, 1956; 1960; 1962;
Soliman and Vondra, 1988; and Chapter II of this report) have considered the Sykes Mountain Formation of the Bighorn
Basin to have been deposited under tidal and shallow marine conditions. However, as far as the writer is aware, there is no detailed work done regarding the paleocurrent analysis of the Sykes Mountain Formation of the Bighorn Basin.
However, some paleocurrent studies have been done on the
Dakota Sandstone of the Ottawa County, Kansas by Franks et 86
al. (1959) and on the Fall River Formation in the Wyoming
Rockies by Mackenzie and Ryan (1962).
This chapter deals with the primary sedimentary structures as paleocurrent indicators and their use in the
determination of the paleocurrent flow directions and the
possible source areas for the sediments of the Sykes
Mountain Formation.
Paleocurrent Measurement
The most significant primary sedimentary structures
which can be used as paleocurrent indicators in the Sykes
Mountain Formation include cross-stratification, ripple
marks, and fossil pelecypod orientation. The cross- stratification includes planar, trough, and herringbone cross-stratification. Ripple marks include both symmetrical
(oscillation) and asymmetrical (current) types. The
majority of paleocurrent measurements were obtained from the
sandstone bodies of the Quartz Arenite Facies Association.
The significance of these primary structures and other
structures present in the Sykes Mountain Formation have been
discussed in Chapter II. At most localities studied, one or
more of these paleocurrent indicators is present.
Cross-stratification and ripple marks correspond very
well to local flow direction, as proven by many studies ' 87
(Potter and Pettijohn, 1977). Several studies of cross-
lamination and ripple marks have proven their importance in
determining the direction of paleoflow, paleoslope, and
location of probable source areas (Potter and Olson, 1954;
Brett, 1955; Tanner, 1955; Potter and Siever, 1956;
Pettijohn, 1957; 1962; Cazeau, 1960; Evans, 1965; Briggs and
Cline, 1967; Klein, 1967; Michelson and Dott, 1973; and many
others). The direction of dip of the planar cross- stratification reflects the direction of sediment transport
or flow of currents which deposited them (Franks et al.,
1959). Directions of the trough cross-stratification axes
provide a good index of paleocurrents (Dott, 1973) and can
be used as an index of sediment transport (Michelson and
Dott, 1973). The majority of ripples with relatively straight crests are commonly transverse to the current direction (Potter and Pettijohn, 1977) with the steeper slope downcurrent. Current-oriented fossils may help in determining the direction of paleocurrent flow. "Elongate
forms that have clearly differentiated ends, ..., probably
indicate a direction of movement, the pointed end usually
points up-current into the direction of flow" (Potter and
Pettijohn, 1977, p. 44). Several current-oriented fossils
are either normal or parallel to current direction
(Seilacher, 1960). 88
In order to determine the direction of the paleocurrent flow in the Sykes Mountain Formation, paleocurrent measurements were made wherever possible at every exposure measured and described in the area of study. Three important paleocurrent indicators were most useful. These were cross-stratification, ripple marks, and current oriented fossil pelecypods. The paleocurrent direction was determined by measuring the azimuth of the cross-bed dip direction of the planar cross-stratification, the bearing of the trough axes of trough cross-stratification, the direction of inclination of the lee side of the asymmetrical ripples, and the direction of elongation of the current- oriented pelecypod fossils, assuming that the pointed end points up-current. The number of paleocurrent measurements obtained from different paleocurrent indicators was variable from one location to another depending on their abundance and preservation. A total of 905 paleocurrent measurements have been obtained; of which 348 measurements were obtained from cross-strata inclination directions measured at 13 locations, 467 from asymmetrical ripple marks measured at 18 locations, and 90 from current oriented fossil pelecypods measured at 4 locations.
Since the Sykes Mountain Formation is no longer horizontal, it was necessary to correct the attitude of each 89 paleocurrent measurement for regional tilt by means of the steriographic net as demonstrated by Ramsay (1961). Tables
4, 5, 6, and 7 are summaries of the different directional structures at different locations.
TABLE 4. Summary of ripple measurements at different locations
Section N MWL MH Dom. Sub. (cm) (mm) dir. dir.
SNA. 2 5 — — 135 15 SNA. 3 23 - - 15 165 SMA. 4 5 - - 135 345 SMA. 7 66 - - 285 105 LSMA. 8 44 10.8 1.4 165 225 LSMA. 9 29 10.6 0.9 135 315 LSMA. 10 15 24.4 2.2 225 345 RD. 11 4 18.0 1.1 345 165 RD. 12 51 13.2 1.4 105 225 RD. 13 69 - - 285 75 SM. 14 9 12.4 2.3 285 105 SM. 15 21 16.3 2.0 315 225 SM. 16 10 19.5 2.5 135 215 RPA. 17 77 12.3 1.5 225 15 HY. 20 9 10.0 1.0 150 0 TH. 21 10 12.3 1.4 285 75 CODY. 22 14 9.8 1.8 105 - CODY. 24 6 13.3 2.0 150 45 1 1 — — Average 1 1 1.6 193 156
N = number of observations.
MWL = mean wave length.
MH = mean height.
Dom. dir. = dominant direction.
Sub. dir. = subsidiary direction. 90
TABLE 5. Summary of cross-stratification measurements
Section N Mean Set Dom. Sub. dip thickness dir. dir. angle (cm)
SMA. 3 4 — — 75 195 SMA. 4 30 17.1 - 285 135 SMA. 7 14 - - 315 255 RD. 11 3 10.0 11.0 165 285 RD. 12 123 16.2 18.9 135 195 RD. 13 15 30.8 - 285 165 SM. 14 30 18.6 30.6 255 330 SM. 15 5 12.0 15.0 255 75 SM. 16 6 18.8 34.0 285 135 RPA. 17 12 14.8 21.9 165 15 TH. 21 32 18.6 11.4 285 135 CODY. 22 27 21.3 19.2 0 75 CODY. 24 47 15.9 37.4 315 135
Average 17.6 22.2 217 164
N = number of observations.
Dom. dir. = dominant direction.
Sub. dir. = subsidiary direction.
TABLE 6. Summary of current oriented fossils measurements
Section Number of Dominant Subsidiary observations direction direction
SMA. 7 30 225 315 LSMA. 9 4 315 195 RD. 11 31 225 - RD. 13 25 195 -
Average 240 127.5 91
TABLE 7. Summary of all paleocurrent measurements at different locations
Section Number of Dominant Subsidiary observations direction direction
SMA. 2 5 135 15 SMA. 3 27 15 165 SMA. 4 35 135 285 SMA. 7 110 285 105 LSMA. 8 44 165 225 LSMA. 9 33 135 315 LSMA. 10 15 225 345 RD. 11 38 225 - RD. 12 174 105 195 RD. 13 109 285 195 SM. 14 39 255 105 SM. 15 26 315 210 SM. 16 16 135 285 RPA. 17 89 225 15 HY. 20 9 150 0 TH. 21 42 285 75 CODY. 22 41 75 0 CODY. 24 53 135 315
Graphical presentation of data
All directional data were graphically represented in the form of 30-degree sector rose diagrams (circular histograms) at the different locations, showing the frequency percentages of azimuths in each sector. The dominant as well as subsidiary paleocurrent flow directions are shown in the above tables. The azimuths of the dominant and subsidiary paleoflow directions were transferred to regional maps to show the paleocurrent directions from all 92 paleocurrent indicators. Figures 21 and 22 are the rose diagrams and the regional distribution of the dominant and subsidiary paleocurrent directions, respectively, at outcrops where paleocurrent measurements of the ripple marks were measured. Figures 23 and 24 show the graphical presentation of the paleocurrent data obtained from the cross-stratification measurements. Figures 25 and 26 for the data obtained from the current-oriented pelecypod fossils, and Figures 27 and 28 for the measurements obtained from all paleocurrent indicators combined at each location.
Discussion
The above tables, rose diagrams, and regional paleocurrent distribution maps show that paleocurrent flow directions from all sources are variable in spite of the dominance of certain directions. Ripple and cross- stratification measurements show that the dominant paleocurrent mostly flowed westwards, and the subsidiary current generally flowed eastwards. However, other paleocurrent directions occur. The current-oriented pelecypod fossils indicated paleocurrent direction mainly to the SW. Most of the rose diagrams of the paleocurrent measurements in the Sykes Mountain Formation are bimodal and some are bipolar. However, some of these circular histograms are unimodal. The rose diagrams based on the 93
SMA 7 SMA 2 66 SMA 4 SMA 3 23
SM1S SM 16 69
LSMA9 39
CODY 23 SM 14 14 CODY 34
TH 31 10 HY 30
RPA 17
RO 13
MAIN DIRECTION SUBSIDIARY DIR.
30 RD12 : LOCATION NO. OF observations 20 30 40%
FIGURE 21. Rose diagrams from ripple measurements 94
109 108® 107 # BILLINGS
BEARTOOTH MTNS /?>v PRYOR MTNS.%^/<^ AVk V; loveu. * # v/{> .*6» SHERIDAN
BIGHORN BASIN
WORLAND 44 — XXXXY %> TENSLEEP XXXXXX ^ YIFU
• THERMOPOLIS
,\W1ND RIVER MTNS
WIND RIVER BASIN
108 107
Dominant Direction Subsidiary Direction
FIGURE 22. Regional distribution of dominant and subsidiary paleocurrent directions from ripple measurements 95
SMA 7 SMA3 SMA 4 30
TH 2} . 32 SM IS
FIGURE 23. Rose diagrams from cross-stratification measurements 96
109 108 107 BILLINGS
30 mi % . >v7v BEARTOOTH MTNS \ PRYOR MTNS.%%)/A MT %Z._ 1" 5®_
"'J/' LOVEU.
SHERIDAN
O :^CODY • SHELL\'1' w < GREYBULL 3
BIGHORN BASIN
g.#C WORLAND 44*-— x*x%% :f xxxxxx TENSLEEPY;:\^-'/>\>3. br, v»-c\ ^^s^HERMOPOLIS ff
,\WIND RIVER MTNS. 'V\- wind river basin
109 108 107 -t Dominant Direction Subsidiary Direction
FIGURE 24. Regional distribution of dominant and subsidiary paleocurrent directions from cross- stratification measurements 97
SMA7 30
LSMA9 RD 13
•MAIN DIRECTION SMA 7 : LOCATION ^SUBSIDIARY DIR.
OBSERVATIONS 30 10 20 30 40%
FIGURE 25. Rose diagrams from current-oriented pelecypod fossil measurements 98
BILLINGS î 30mi
BEARTOOTH MTNS. ,PRYOR MTNS.^>VC
lOVELL I / 1 SHERIDAN
CODY
GREYBULL
BIGHORN BASIN
WORLAND « XXXXY :i % Wi xxxxxx fp > TENSLEEP x/xx ^4%:; hfàî # THERMOPOLIS
,\WIND RIVER MTNS.
WIND RIVER BASIN
_ _C
Dominant Direction Subsidiary Direction
FIGURE 26. Regional distribution of dominant and subsidiary paleocurrent directions from current-oriented pelecypod fossil measurements 99
SMA 7 1 "O SMA 4 SMA 3 ^ 3» 27
ISMA 6 44
38
RD 12 174
ISMA 9 , 33 ISMA lo RD 13 109
SMI4 39
HY20 RPA 17 SM16 89 _
MAIN DIRECTION TH 21 42 SUBSIDIARY DIR
30
TH 21 I LOCATION CODY24 42 t NO. OF OBSERVATIONS S3
FIGURE 27. Rose diagrams from measurements obtained from all indicators combined 100
109 108" 107 BILLINGS
30mi
BEARTOOTH MTNS PRYOR MTNS.'> ""
SHERIDAN
CODY
GREYBULL 3 BIGHORN BASIN ?Êâ\
WORLAND 44^—— TENSLEEP
1#\ \ , m THERMOPOLIS tii'- ,\WIND RIVER MTNS. •'-VV.Vvv-V WIND RIVER BASIN
108 107
Dominant Direction Subsidiary Direction
FIGURE 28. Regional distribution of dominant and subsidiary paleocurrent directions from all indicators combined 101 measurements of the ripple marks and the cross- stratification are less consistent than those based on fossil measurements. The presence of bimodal and bipolar histograms are due to the occurrence of two different paleoflow directions reflecting the to and fro movement of tidal currents during the deposition of the Sykes Mountain
Formation. Such bidirectional current flow is a characteristic of tidal flat deposits. The bimodal-bipolar distributions are a result of ebb and flood currents (Swett et al., 1971). Stanley and Swift (1976) and Kreisa (1981) reported that shallow marine shelves are characterized by diversely trending currents that are generated by storms.
So, this may explain the variable distribution of the paleocurrents recorded by the Sykes Mountain Formation since its environment of deposition was dominated by tidal currents and periodic storm activity. Moreover, this kind of distribution of paleocurrents could be due to rotary tidal flow (Banks, 1973). Bimodal dispersal patterns characterizes beach sediments (Klein, 1967). Tidal flats are among several environments associated with coastal sedimentation which are characterized by two current systems; the ebb and flood (Klein, 1967). In the Sykes
Mountain Formation, the dominant paleocurrent direction represents the ebb current and the subsidiary paleocurrent 102 represents the flood current. Therefore, in this particular case, the ebb tidal current velocities exceeded the flood current velocities and reflected the probable location of the source area.
As mentioned above, the sedimentary structures help in the determination of the direction of sediment transport and hence the probable location of source areas. Regarding the relation of the paleocurrent direction and the source area,
Pettijohn (1962, p. 1484) states: "The source area is upcurrent. Conceivably, if the deposits are marine, the current might be a littoral current and the source elsewhere than upcurrent. But a widespread uniform current pattern can hardly be due to littoral currents and the upcurrent direction must also be the source direction." All the paleocurrent indicators used for the determination of the paleocurrents in the present study indicate that the general dominant paleocurrent flow direction (ebb current) during the deposition of the Sykes Mountain Formation sediments was to the southwest, and the general subsidiary direction
(flood current) was to the southeast. According to
Pettijohn (1962), usually (but not always), the grain size decreases in a downcurrent direction. This may explain the decrease in grain size of the Sykes Mountain Formation deposits from the west (along Sheep Mountain Anticline) to 103 the east (at Beaver Creek area). The paleocurrent regional distribution maps show that there are generally two dominant directions of paleocurrent flow —easterly as well as westerly directions— indicating that transportation of the
Sykes Mountain Formation sediments in the Bighorn Basin was essentially from the east as well as from the west.
Therefore, the probable main source of the Sykes Mountain
Formation sediments was to the east and west of the study area. These inferences of the source areas will be further documented in Chapter IV. Working on cross-stratification in the Dakota Sandstone of Ottawa County, Kansas, Franks et al. (1959) reported an average paleocurrent direction of
S57W and concluded that the currents which were responsible for the transportation of the Dakota Sandstone sediments flowed mainly from the northeast to the southwest and hence the source of the Dakota sediments in Ottawa County was to the northeast. Moreover, Mackenzie and Ryan (1962, p. 49) reported that northwestward and westward flowing currents transported the sediments of the Fall River Formation and its equivalents in the Black Hills and in Central Wyoming.
They added that sediments at the base of the Sykes Mountain
Formation in the Owl Creek Mountains were transported toward the northeast and east. Although he did not measure the paleocurrents in the Sykes Mountain Formation in detail. 104
Moberly (1960) reported that foresets of the cross- stratified sandstones near the base dip west or west- southwest, indicating westward sediment transport. The results of the present study agree with Moberly's (1960) observations. The paleocurrent measurements of the present study show that the sediments of the Sykes Mountain
Formation of Bighorn Basin could have been deposited by currents related to the currents of the Craton dispersal system, the paleocurrents of which flowed northwest, west, and southwest across the study area and distributed sediments generally westward (Mackenzie and Ryan, 1962).
Moreover, from their pétrographie studies, Mackenzie and
Ryan (1962) reported that the sediments deposited by paleocurrents related to the Craton dispersal system were derived from a land mass of mainly sedimentary and metamorphic rocks located in the region of the Sioux Arch and the southwestern Canadian Shield to the east of the study area. In addition, some paleocurrent measurements of the present study show current flow to the northeast, east, and southeast indicating the presence of a source area to the southwest and northwest of the study area which may contribute some detritus eastward from the geanticline into the sedimentary basin. These currents may relate to the
Idaho dispersal system of Mackenzie and Ryan (1962). The 105 contribution of detritus from the western source is much less than that of the eastern source since the paleocurrent flow direction is generally to the west and southwest.
Summary
Paleocurrent indicators in the Sykes Mountain Formation include ripple marks; planar, trough, and herringbone cross- stratification; and current-oriented pelecypod fossils.
Paleocurrent flow direction was determined by measuring the azimuth of the direction of the maximum inclination of the planar cross-stratification, the direction of trough axes of the trough cross-stratification, the direction of inclination of the steeper lee slope of the asymmetrical ripples, and the direction of elongation of the current- oriented pelecypod fossils, with the pointed end pointing upcurrent. The measured directions were corrected for the regional tilt using the steriographic net. Rose diagrams for different paleocurrent indicators were constructed and regional distribution maps for the dominant as well as the subsidiary paleocurrent direction were prepared.
Paleocurrent flow directions lie almost in all directions with dominant easterly and westerly directions, indicating that the currents which transported the sediments of the Sykes Mountain Formation flowed mainly from the east 106 and west and suggesting that the probable source of the sediments was located to the east and west of the study area. 107
CHAPTER IV. PETROLOGY
Introduction
Many pétrographie studies eoneerning the Sykes Mountain
Formation and its equivalents have been earried out by
several authors. Moberly (1956; 1960) studied the petrology
of Morrison, Cloverly, and Sykes Mountain formations in the
Bighorn Basin, Wyoming and Montana. MacKenzie and Poole
(1962) diseussed the provenance of Dakota Group sandstones
of the Western Interior. MaeKenzie and Ryan (1962) earried
out pétrographie analyses on the Cloverly-Lakota and Fall
River sandstones in the Wyoming rockies. Chisholm (1963)
made extensive studies of the heavy minerals of the Upper
Jurassic and Lower Cretaceous sediments of the Western
Interior. Kvale (1986) analyzed some mudstone samples which
were provided by the writer from the Sykes Mountain
Formation.
This chapter of the dissertation is a discussion of the
petrology of Sykes Mountain Formation based on the writer's
field and thin section observations as well as the data
obtained by various previous workers. 108
Petrology of Sandstones
Sandstones of the Sykes Mountain Formation are well indurated, fine- to very fine-grained quartz arenites. The visual comparators provided by Rittenhouse (1943), Powers
(1953), and Beard and Weyl (1973) were used for estimating sphericity, roundness, and grain size and sorting of the sand grains, respectively, because of the inaccuracy of sieve analysis of indurated sandstones, especially those which are compacted and silica cemented (Beard and Weyl,
1973). Moreover, measurements of grain-diameter distributions from thin sections are not precise because this technique measures grain intercepts along traverse lines which do not intersect all grains at their maximum thin-section diameters (Beard and Weyl, 1973). Sandstones of the Sykes Mountain Formation are fine- to very fine grained, with detrital grains ranging from subangular to subrounded, along with the occurrence of a few angular and rounded grains. The sphericity of detrital grains ranges from 0.73 to 0.83. These sandstones are usually well sorted.
Quartz is the chief detrital component. The quartz grains are subangular to subrounded. However, a few rounded grains occur. Accessory minerals are feldspars (orthoclase, microcline, and plagioclase), muscovite, and heavy minerals. 109
Monocrystalline as well as polycrystalline quartz grains with both uniform and undulatory extinction are present.
Some of these quartz grains show quartz overgrowths.
Stratification is clear in some samples, and the long dimensions of many elongated grains are parallel to the bedding.
Grain contacts
The grain-to-grain contacts present in the sandstones of the Sykes Mountain Formation include concavo-convex, sutured, and long. The long contacts could be developed as a result of original packing of straight-edged sand grains, secondary cement precipitation, or pressure (Taylor, 1950).
Concavo-convex grain contacts may be due to either solid flow of the yielding grain, or solution at points of grain contacts and deposition or removal of the dissolved material. Some of the concavo-convex contacts could be due to original packing and grain shape. Interlocking of the quartz grains is well developed in the sandstones of the
Sykes Mountain Formation and could be the result of solution of silica at grain points of contacts and precipitation of secondary silica. Sutured contacts are serrated to wavy lines of contact which may be produced from local solution.
Microstylolites in sandstones similar to those described by many authors (e.g., Sloss and Feray, 1948; Heald, 1955) 110 could have resulted from the development of this type of contact by increasing pressure and differential solution.
Cementation
The quartz arenites of the Sykes Mountain Formation are siliceous (cemented by very finely crystalline silica), calcareous, and/or ferruginous. The dominant cementing materials are secondary quartz, calcite, or iron oxides.
One or more of these cements may occur in the same sample.
Sources of secondary silica Many authors have discussed and debated the origin of secondary silica which is precipitated as quartz cement in sandstones. Most of the silica cement in sandstone probably originated from the intergranular pressure solution of sand grains at their points of contact (Waldschmidt, 1941; Waugh, 1970; Pittman,
1972; Levandowski et al., 1973; Blatt, 1979; 1982). The dissolved silica is either precipitated immediately or may be carried in solution to be precipitated elsewhere. In the sandstones of the Sykes Mountain Formation, the presence of pressure solution is indicated by the occurrence of interlocked quartz grains and sutured grain contacts. In addition to the intergranular pressure solution process as the most important mechanism of originating secondary silica, there are other possible sources; 1) silica released during clay mineral diagenesis in adjacent shales Ill
(Towe, 1962; Blatt, 1979; Boles and Franks, 1979; Schmidt
and McDonald, 1979), 2) decomposition of feldspars
(Fothergil, 1955; Schmidt and McDonald, 1979), 3)
decomposition of silica tests secreted by organisms, such as
those of diatoms and radiolaria (Siever, 1957; Von
Englehardt, 1967), 4) silica released during replacement of
quartz grains by carbonate (Walker, 1960; 1962; Levandowski
et al., 1973; Schmidt and McDonald, 1979), and 5)
precipitation from meteoric water and/or connate water
(Levandowski et al., 1973), or directly from sea water
(MacKenzie and Gees, 1971; Levandowski et , 1973). It is
very difficult to determine which of these silica sources
might be more important than that of pressure solution.
Sources of carbonate cement Quartz arenites of the
Sykes Mountain Formation contain significant amounts of
calcium carbonate cement. The probable sources of the
carbonate cement can be explained as follows; 1)
Dissolution of carbonate fragments such as shell fragments
and limestone grains due to the action of both the pressure
from overburden and temperature change (Waldschmidt, 1941;
Blatt, 1979). The scarcity or absence of fossil remains in
the Sykes Mountain Formation may support the above
explanation as the source of calcite in its sandstones. 2)
Clay mineral diagenesis yields not only silica but also ions 112 of calcium, magnesium, and iron for sandstone cementation
(Boles and Franks, 1979). 3) According to Waldschmidt
(1941, p. 1865), precipitation of calcite within the pore spaces of the sandstones can occur when pressure is released and temperature is increased. Siever (1959) noted that calcite solubility decreases and precipitation takes place, while silica solubility increases, as the temperature increases during the deep burial of sandstones. 4) Calcium carbonate could also be precipitated from both meteoric water and marine water (Blatt, 1979). According to several authors, calcareous cementation of sandstone occurs at very shallow depths (e.g., Todd, 1963; Warner, 1965) and is usually related to the presence of unconformities (Blatt,
1979).
Origin of iron oxide cement Iron oxide was precipitated as coatings of the detrital grains and as pore filling. The main iron oxide cement in the sandstones of the Sykes Mountain Formation is hematite. The formation of hematite requires a source of iron atoms and circulation of meteoric water to maintain oxidizing conditions in the pore spaces (Blatt, 1979). The iron atoms necessary for hematite formation are mainly derived from the weathering of iron- rich silicates (Boles and Franks, 1979; Blatt, 1982). 113
Heavy minerals
The thin section study of sandstones from the Sykes
Mountain Formation revealed the occurrence of different heavy minerals representing less than 1% of the bulk mineral composition and including abundant opaque minerals. The nonopaque heavy minerals (beginning with the most abundant) include zircon, tourmaline, garnet, rutile, staurolite, chlorite, and biotite. No quantitative analysis of the heavy minerals present has been carried out by the writer.
The previous workers' data (e.g., Moberly, 1956; 1960;
MacKenzie and Ryan, 1962; MacKenzie and Poole, 1962;
Chisholm, 1963; 1970) regarding the heavy minerals in the sandstones of the Sykes Mountain Formation should be sufficient for this study.
In his study of the Morrison, Cloverly, and Sykes
Mountain formations in the Bighorn Basin, Moberly (1956,
Fig. 10; 1960, Table 1) reported the common occurrence of leucoxene and black opaque minerals; the presence of zircon, chlorite, and rutile; and the rare occurrence of garnet, tourmaline, and authigenic pyrite in the sandstones of the
Sykes Mountain Formation. He (Moberly) added that the amounts of leucoxene and black opaques are equal, the tourmaline grains are either angular or rounded and euhedral and subangular to subrounded zircon grains are present. 114
MacKenzie and Ryan (1962) have studied the petrography
of the Fall River Formation and its equivalents in the
Wyoming Rocky Mountains. They reported that the sandstones
are composed of rounded to subrounded, and sometimes
angular, quartz grains with uniform and undulatory
extinction, polycrystalline silica clasts, Muscovite, and
rare plagioclase feldspar, and they are cemented by silica,
carbonate, and iron oxides. Furthermore, they added that
75% of the average heavy mineral suite is represented by
rounded zircon and tourmaline (with more tourmaline than
zircon) and the rest includes staurolite, subangular to subrounded garnet, chloritoid, kyanite, epidote, and rare
rutile.
MacKenzie and Poole (1962), on the basis of their
mineralogical study of the Dakota sandstones of the Western
Interior, recognized a western suite and an eastern suite of detrital minerals. The western suite is characterized by
the dominance of rounded tourmaline, presence of chert, and
absence of muscovite, feldspar, and chloritoid. The eastern suite is characterized by dominant angular tourmaline, presence of feldspar, muscovite, and chloritoid and rare chert. They added that the sandstone units above the
regional disconformity (Sykes Mountain Formation) contain
the detrital minerals characterizing the western suite in 115
the western part of the area and those characterizing the
eastern suite in the eastern part.
Chisholm (1963) studied the petrology of the Upper
Jurassic and Lower Cretaceous strata of the Western
Interior. He recognized nine distinctive heavy mineral
assemblages; zircon-tourmaline suite of the Idaho-Wyoming
thrust belt, chloritoid-volcanic suite of the Muddy
Formation, garnet suite of the Muddy Formation, zircon-
tourmaline suite of the Colorado Front Range, chloritoid
suite, staurolite suite, zircon-tourmaline suite (of the
Dakota), garnet suite, and zircon-tourmaline suite of the
Lower Morrison Formation. The most obvious feature of the
characteristics of these nine heavy mineral suites is that
zircon and tourmaline represent the majority of all heavy
mineral fractions (40-95%) with variable contents of other
minerals.
According to Chisholm (1963), the sandstones of the
Sykes Mountain Formation are characterized by three out of
these nine heavy mineral assemblages: 1) the staurolite suite, 2) the chloritoid suite, and 3) the zircon-tourmaline suite. Figure 29 shows a typical heavy mineral composition
of the Sykes Mountain Formation.
The staurolite heavy mineral suite mainly contains zircon and tourmaline, with variable amounts of other heavy 116
ANGUI^
o si TOURMAUNE I
0-044 — 0.125 MM 0.044-0.125MM
Itf
0.125 - 0.250 MM A ANGULAR CH CHLORITE R ROUNDED XLS CRYSTALS
FIGURE 29. Typical heavy mineral composition of the Sykes Mountain Formation (after Chisholm, 1963) 117 minerals, including staurolite, rutile, garnet, diaspore and kyanite, with chlorite absent, whereas the chloritoid suite is essentially composed of zircon and tourmaline, with abundant chloritoid, chlorite, rutile, garnet, common apatite, and staurolite, kyanite, and diaspore are absent
(Chisholm, 1963; 1970). Chisholm (1963) added that the chloritoid-bearing sandstones are generally finer-grained than staurolite bearing sandstones and rounded zircon and tourmaline grains in the staurolite suite are more common than in the chloritoid suite. "The zircon-tourmaline bearing sandstones are very similar to the Lakota or Front
Range type sandstones" (Chisholm, 1963, p. 76). Figure 30 shows the distribution of the heavy mineral suites of Sykes
Mountain Formation and Upper Lakota.
Provenance
The main objective of the mineralogical study of the
Sykes Mountain Formation is to determine the lithology of its source area. According to Moberly (1956; 1960), the heavy minerals of these sediments have been derived from pre-existing sedimentary rocks, with occasional small additions from metamorphic and pegmatitic rocks or from acid igneous rocks, but he did not determine the ultimate source of these sediments. MacKenzie and Ryan (1962) reported that the Fall River sandstone and its equivalents are related to 118
montana a utah'P ^.^.SJ!A5 K A
c OL ORAdO
staurolite sandstone logsl zircon —tourmaline 1 OOO sandstone of idaho- mm chloritoid sandstone wyoming thrust belt
jsga mixed chloritoid— percentage of rounded l^kal staurolite sandstone S zircon
FIGURE 30. Distribution of the heavy mineral suites of Sykes Mountain Formation and Upper Lakota (after Chisholm, 1963) 119 the cratonic dispersal system and were derived from a source area of sedimentary and metamorphic rocks in the region of
Sioux Arch and the southwestern Canadian Shield east of the study area. The eastern and western suites recognized by
MacKenzie and Poole (1962) were derived from the Precambrian rocks of the Canadian Shield and the pre-existing sediments of the Cordilleran region to the west, respectively.
Sandstones of the staurolite suite recognized by Chisholm
(1963) were derived from an area of medium-grade metamorphic rocks with pre-existing sedimentary rocks. According to
Potter and Pryor (1961), neither the Precambrian rocks of the Canadian Shield, nor the Paleozoic rocks of the mid- continent area could be the source of the staurolite-suite sandstones. The minerals found in the staurolite-bearing sandstones (Chisholm, 1963) are similar to those found by
Groot (1955) in the Lower Cretaceous of Delaware and the adjacent area. The source area of the staurolite-bearing sandstones must have appreciable amounts of metamorphic rocks and hence is likely the Piedmont Province of the
Appalachian area (Chisholm, 1963). The minerals of the chloritoid suite (Chisholm, 1963) must have been derived from low-grade crystalline schists (chloritoid schists and chlorite schists). Moreover, volcanic and pre-existing sediments contributed variable amounts of material to the 120 chloritoid sandstones. It is likely that the chloritoid suite was derived from a source area to the north, along the borders of the Canadian Shield and that the tourmaline- zircon suite is derived from pre-existing sedimentary rocks
(Chisholm, 1963).
Mudrocks
As pointed out above, the Sykes Mountain Formation is composed essentially of quartz arenite, siltstone, and dark gray and black mudstones and shales. The entire formation weathers rusty brown to yellowish brown. The units of intercalated sandstones, siltstones, and shales have been interpreted as intertidal and subtidal deposits (Chapter
II). Plant stem fragments, plant leaf parts, and carbonaceous materials are common within the mudstone units.
Ironstone laminae, nodules, and concretions are also abundant. Pyrite nodules occur also but are less common.
At some places, cone-in-cone concretions occur at the base of some mudstone units. At the base of the overlying marine black Thermopolis Shale and above its contact with the Sykes
Mountain Formation, a bed of dahllite concretions (calcium phosphate) occurs. The presence of plant remains, carbonaceous material, and pyrite in Sykes Mountain
Formation indicate that reducing conditions may have 121 occurred at some time after deposition (Moberly, 1960). The main sedimentary structure in these mudrocks is the fissility which is due to the abundance of clay minerals
(Blatt, 1982) and later load compaction of the sediment
(Moberly, 1956).
According to Moberly {1956, Fig. 70; 1960, Fig. 11, p.
1168), clay mineral assemblages of the Sykes Mountain
Formation contain mixtures of kaolinite, chlorite, montmorillonite, and illite, with an abundance of kaolinite.
Also, Kvale (1986, Table 2 and Appendix C) analyzed some mudstone samples from the Sykes Mountain Formation and reported that the formation contains varying amounts of kaolinite, mixed-layer montmorillonite-illite, and illite.
He interpreted this suite of clay minerals as having been deposited on a shallow marine shelf.
The clay mineral types and composition of the mudrocks are controlled by source rock composition, climate, relief, tectonism, and diagenetic processes (Biscaye, 1965; Blatt et
al., 1972; Weaver, 1978; Singer, 1980; 1984). According to
Weaver (1967), illite and chlorite are the most abundant
clay minerals in Paleozoic mudrocks, while kaolinite and
mixed-layer clay minerals are dominant in younger mudrocks.
These trends of clay mineral changes through time are a
result of the alteration of mixed-layer clays to illite and 122 chlorite during diagenesis (Hiltabrand et , 1973; Hower et al., 1976). According to Bhatia (1985), mudrocks that are rich in phyllosilicates and poor in feldspars suggest a multicycled nature and intense weathering conditions in the source area. Moreover, he has attributed the occurrence of chlorite in mudrocks either to older rocks rich in chlorite or to volcanic smectite. Weaver (1978) reported that diagenetic changes in marine conditions are minor; and the majority of chlorite and illite are attributed to a terrigenous origin. On the contrary, Blatt et (1972) attributed the formation of illite to the diagenetic processes related to burial.
Dahllite concretions
It has been mentioned earlier that there is a thin bed of shale rich in dahllite concretions at the Sykes Mountain
Formation-Thermopolis Shale boundary, which has been chosen as a marker of that boundary. These concretions have also been reported by several other workers in the Bighorn Basin
(e.g., McConnell, 1935; Eicher, 1960; 1962). McConnell
(1935) reported them in the region near Cody. Other occurrences were reported elsewhere in areas other than the
Bighorn Basin. The concretions described by Rogers (1912) were found in the Utschina River area, Podolia, USSR. Small radial pellets of dahllite were found by Washington (1929) 123 on St. Paul's Rocks. The concretions from Wyoming are smaller than those from Russia and larger than those from
St. Paul's Rocks, and more structured than both.
The dahllite concretions from the area of study are spherical. They range in diameter from 1 to 4 cm or more and have a rough surface. Internally they possess a radial structure. They are unusual because of their structure and composition (McConnell, 1935). Figure 31 shows the nature of their external surface and their internal structure. The structure and origin of the mineral dahllite (carbonate- apatite) has been controversial (McConnell, 1960).
McConnell (1935) reported that the concretions were formed in situ in the shale and are not the result of replacement or "the concretions may represent replacement of some other material but, if this be the case, the replacement was quite complete, because no relicts of the replaced material were found."
According to Hamilton (1973), occurrence of large dahllite concretions in nonphosphatic shales reflects unusual conditions at their time of formation. To determine the conditions responsible for the formation of the dahllite concretions, Hamilton (1973) attempted to synthesize phosphatic nodules of similar composition and structure by the gel growth technique. Although Hamilton (1973) expected 124 m
FIGURE 31. Dahllite concretions from outside and inside
carbonate-apatite (dahllite) to be formed, he found that silicate-phosphate spherulites had formed instead. He attributed the strong similarity in the structures of the silicate-phosphates obtained and the dahllite concretions in question to possible similarity of growth process.
Moreover, he concluded that "natural phosphate nodules formed by the progressive phosphatization of initial poorly crystallized silicate-phosphate nodules. Nodules formed by this process have indeed formed _in situ, but not as the result of direct crystallization but rather as the result of diagenesis." 125
Furthermore, a dahllite-apatite-rich horizon in the shale of the Middle Devonian Moscow Formation of New York has been used by Baird (1978) for the recognition of a submarine discontinuity. According to Baird (1978), most of the nodules appear to have been formed in near-surface muds after deposition from interstitial precipitation of calcium phosphate.
Summary
Sandstones of the Sykes Mountain Formation are fine- to very fine-grained quartz arenites, with subangular to subrounded detrital grains. Mineralogically, they are mainly composed of quartz grains with accessory feldspars, muscovite, and heavy minerals. Quartz grain-to-grain contacts include concavo-convex, sutured, and long types.
Dominant cementing materials are secondary quartz, calcite, and iron oxides with the possibility of the presence of one or more of them in the same sample. Heavy minerals include zircon, tourmaline, chloritoid, staurolite, rutile, diaspore, chlorite, kyanite, garnet, and others. The sediments of the Sykes Mountain Formation were likely derived from metamorphic rocks of the Canadian Shield.
Contribution of detritus was derived from the north along the borders of the Canadian Shield, and from the sediments 126 of the Cordilleran region to the west. Clay minerals include abundant kaolinite with variable amounts of other clays (illite, mixed-layer montmorillonite-illite, and chlorite). The dahllite concretions which occur at the
Sykes Mountain Formation-Thermopolis Shale boundary were formed iri situ in the shale as a result of diagenesis. 127
CHAPTER V. TECTONICS AND SEDIMENTATION OF LOWER CRETACEOUS
SEDIMENTARY ROCKS
Tectonics
The Bighorn Basin is among those basins within the
Wyoming Province which extends from the western edge of the
central Great Plains to the Idaho-Wyoming thrust belt and
from southwestern Montana to northern New Mexico (Figure
32). Several authors reported that these basins and uplifts
were formed as a result of partitioning of the early Sevier
foreland basin by the Laramide Orogeny.
The Sevier and Laramide orogenies are the main tectonic
movements that have affected the Rocky Mountain region. The
tectonism began in the overthrust belt to the west during
the Late Jurassic and in the foreland on the east during the
Late Cretaceous, with an overlap in time of movement in the
two regions after the middle Late Cretaceous (Wiltschko and
Dorr, 1983). According to Armstrong and Oriel (1965), the
early phase of tectonism occurred in the miogeosyncline.
They reported the occurrence of a northeast-trending uplift
in Idaho in Late Triassic time reflecting "the start of the
breakup of the miogeosyncline." During Middle and/or Late
Jurassic time, an uplift associated with an episode of overthrusting occurred within the Idaho-Wyoming overthrust 128
I Big Snowy porcupine 110« Upim Domo #BuM# ^ t Mountain.
ND 46°-
«^.44° -/^shuklo Upllil Owl Crook Upiiii
Shirley Mou Uplift
"""
Groat DividtA-g Hannu Basin 42°- Springs Upim C-T^ Laramie H ^'ori-a Madro ^ Basin Uplift 3 , North-Park Sand Wash Basin basin •n
Middle Park a •asin 4
FIGURE 32. Map of portion of Wyoming province showing Laramide sedimentary basins and uplifts (modified from Dickinson et al., 1988) 129
belt (Oriel and Piatt, 1979; Allmendinger and Jordan, 1981;
Wiltschko and Dorr, 1983). Several authors (e.g.. Love,
1960; Armstrong and Oriel, 1965; Oriel and Armstrong, 1966)
reported that thrusting started in the Idaho-Wyoming thrust
belt in the west in the latest Jurassic and ended in the
east in the early Eocene time, with the ages of thrust
faults decreasing to the east. According to Wiltschko and
Dorr (1983), the part of the overthrust belt in the southeast Idaho, northwest Utah, and west Wyoming is broad,
prominent, and convex to the east. The region is cut by some major as well as several minor thrust faults. The
major thrusts from west to east include the Paris, Meade,
Crawford, Absaroka, Darby, and Prospect (Figure 33). These
faults generally trend north-south.
Since foreland basins are formed on cratons as a
characteristic companion of the adjacent mountain ranges,
they represent the major basins of nonmarine sedimentation
(Dickinson, 1974). In addition, foreland basin sediments
reflect the history of deformation of the adjacent mountain
belts (Jordan, 1981). According to Kauffman (1977), the
foreland basin of Idaho, Wyoming, and Utah represents the
western portion of the Cretaceous seaway (Cretaceous Western
Interior basin) that extended from northern Canada to the
Gulf of Mexico, with local width of more than 1,600 km.
Dorr et aj.. (1977) and Wiltschko and Dorr (1983) reported 130
112 110'
Jackson
Ca [GH
43 HOBACK 43 — BASIN
Auburn
on
# La Barge
42 42 — UT
/ WY 41 — /-/ ~ UT" 30mi -I-» 40 km
Pa - Paris Thrust M m Meade Thrust Cr • Crawford Thrust T • Tunp Thrust Pr m Prospect Thrust A » Abaaroka Thrust GH > Qama Hill Fault D • Darby Thrust Ca m cache Fault 112° 111 no'
FIGURE 33. Major faults in the western overthrust belt (from Wiltschko and Dorr, 1983) 131 that the Wyoming part of the Rocky Mountain foreland basin was dominated by episodes of basement-thrusting which were
in large part coeval with the deformation in the Overthrust
Belt during the Late Cretaceous, late Paleocene, and early
Eocene times. Characteristics of the periods of thrusting
and those of little or no thrust displacement have been discussed by several authors (e.g., Beck, 1985; Beck et al.,
1988; Wiltschko and Dorr, 1983; and many others).
Many authors indicated that Laramide structures in northern Wyoming and southwestern Montana deformed the foreland basin and influenced sedimentation patterns, and sometimes have been source areas for the intraforeland basin
(DeCelles, 1984; 1986; Schwartz, 1983; Suttner et al., 1981;
Kvale and Beck, 1985). According to Wiltschko and Dorr
(1983), the Sevier overthrust uplifts mainly involved
Paleozoic and Mesozoic sedimentary sequences, but the
Laramide uplifts are cored by Precambrian crystalline rocks.
The age relationships of the Laramide structures are more complicated than those of the Sevier thrust belt, but a younging-eastward trend occurs (Gries, 1983; Beck, 1985).
The age range of Laramide structures is between 70 and 45 m. y. ago (Coney, 1972), but the maximum deformation of
Laramide Orogeny occurred between 65 and 50 m. y. ago (Berg,
1962). Dickinson et (1988) reported the beginning of 132 the Laramide deformation during Maastrichtian time (65-75
Ma). They added that Laramide deformation ended in the north between early and middle Eocene time (50-55 Ma) and in the south by the end of Eocene time (35-40 Ma).
Dickinson et aJL. (1988) reported that common structural relief between the uplift crests and the basin floors throughout the central Rocky Mountain region is in the range of 5,000 to 10,000 m. However, local structural relief between the uplifts and the intermontane basins in the
Bighorn region has been reported as about 30,000 to 10606 meters and has been attributed to the Laramide Orogeny's compressive forces by Fanshawe (1971), Prucha et (1965), and Stearns (1971; 1975). The uplifts resulting from the
Laramide Orogeny in Wyoming form a broad, asymmetric, arcuate pattern (Figure 34) bounded by zones of thrusting or steep dip and partly encircled deep-basin areas (Berg,
1981). Besides the general northwest-southeast trend of the
Laramide uplifts and basins, a north-south trend occurs in southern Wyoming (Berg, 1981). Moreover, east-west trending
Laramide structures occur, such as the Owl Creek and the
Granite Mountains in Wyoming and the Uinta Mountains in
Utah. The coeval deformation in the Overthrust Belt and the
Wyoming region of the Rocky Mountain Foreland from Late
Cretaceous to early Eocene times (Armstrong, 1968; Dorr et 133 al., 1977; Wiltschko and Dorr, 1983; Beck, 1985; Beck et al.. 1988) indicates that the terms Sevier and Laramide refer to different styles of deformation rather than distinct periods of deformation. Moreover, an episode of
Laramide-style deformation has been reported in the Montana-
Wyoming region of the Rocky Mountain Foreland during the
Early Cretaceous (Schwartz, 1983; DeCelles, 1984; 1986;
Kvale, 1986), indicating that the difference between the
Sevier and Laramide orogenies is one of style only. There has been controversy among geologists regarding the origin and mechanisms of these structures which is beyond the scope of this study.
Sedimentation
According to Beaumont (1978; 1981), Schedl and
Wiltschko (1980), and Jordan (1981), thrusting is an important mechanism for the formation as well as evolution of sedimentary basins. Armstrong and Oriel (1965) related the events in the thrust belt to those in the associated sedimentary basin to the east, which is called the foredeep, foreland, or the Cretaceous seaway. According to Wiltschko and Dorr (1983), the thrusts moving eastward supply a large load consisting of the thrusts and their erosional products.
This load downwarps the lithosphere forming a basin which is 134
MONT YEUowsro WYO N
COLO UINTA 100 mi lOOkm
FIGURE 34. Diagrammatic map of the Wyoming foreland showing arcuate uplifts bounded by zones of thrusting or steep dip (thick curved lines) and deep basin areas (stippled), from Berg (1981) 135 asymmetrical toward the craton. Its deepest part is close to the thrust margin, and its axis moves toward the craton following the thrust load movement direction (Wiltschko and
Dorr, 1983, p. 1305). Episodes of thrusting and uplift are reflected by coarse clastic sedimentation immediately adjacent to the thrust belt, whereas episodes of little or no thrust movements are reflected by decelerations or cessations of coarse clastic sedimentation (Wiltschko and
Dorr, 1983, p. 1305). However, in their recent studies on syntectonic sedimentary facies of Laramide basins. Beck
(1985) and Beck et (1988) reported that coarse clastics were not restricted to the episodes of rapid thrusting and displacement, but they were derived also from thrust tips during periods of slow thrusting.
According to Wiltschko and Dorr (1983), in the latest
Jurassic-earliest Cretaceous, the initial movement of the
Paris and Willard thrusts produced an uplift which was accompanied by the accumulation of the synorogenic, conglomeratic, upper part of the Ephraim Formation adjacent to the thrust belt. The upper Ephraim conglomerate was followed by many alternations of lacustrine or marine and fluvial sediments (Eyer, 1969), reflecting several episodes of thrusting, during which thick fan deposits of coarse detritus (conglomerates and conglomeratic sandstones) were 136 deposited in narrow bands adjacent to the thrust belt, and periods of tectonic quiescence during which coarse clastics eroded from the uplifted rocks were transported to the foreland basin. Beck (1985) and Beck et al. (1988) suggested a facias model for syntectonic sedimentation adjacent to Laramide basement-thrusts. Regarding the sedimentary facies distribution in the Laramide structural basins, they stated; "Characteristic syntectonic sedimentary facies in these basins include a narrow conglomerate facies adjacent to the thrust, a narrow sandstone/mudstone/coal facies immediately basinward, a basinal mudstone/coal/carbonate/evaporite facies along the depositional axis and finally a wide distal sandstone/ mudstone/coal facies." These changes in the facies distribution are possibly due to the variable rates of thrusting and consequent subsidence. The Lower Cretaceous
Gannett Group sequence adjacent to the thrust belt (Eyer,
1969) consists of the fluviatile upper Ephraim conglomerate, lacustrine Peterson Limestone, the fluviatile Bechler
Formation (including conglomerate on the west), lacustrine
Draney Limestone, and marginal lacustrine Smoot formation.
Simultaneous with the deposition of the Upper Ephraim
Conglomerate and the Peterson Limestone adjacent to the thrust belt, sediments were transported to the rest of the 137 foreland basin by streams from eastern and western highlands, as indicated by the paleocurrent measurements made by MacKenzie and Ryan (1962), Mirsky (1962b), Moberly
(1960), and Kvale (1986). The resulting deposits form the lover to middle part of the Lakota Formation in the Powder
River Basin and lower to middle Little Sheep Mudstone in the
Bighorn Basin (Kvale, 1986). The channels of the Lower
Himes Member (Cloverly Formation) in the Bighorn Basin and the fine-grained mudstones of the upper Lakota in eastern
Wyoming were deposited contemporaneously with the deposition of the fluvial Bechler Formation adjacent to the fold-thrust belt during Middle Aptian time. Renewed rapid movement on the Paris Thrust during Late Aptian-Barly Albian time caused fluvial channels to flow westward, resulting in the deposition of the mostly fluvial Greybull interval of the
Himes Member (uppermost Cloverly Formation) in the Bighorn
Basin and the uppermost Lakota fluvial sandstones of Gott et al. (1974) contemporaneously with the deposition of the lacustrine Draney Limestone adjacent to the fold-thrust belt.
By the beginning of Albian time, increased subsidence of the foreland basin resulted in the transgression of the
Early Cretaceous seaway extending from the Arctic Ocean to the Gulf of Mexico (Young, 1970). The transitional marine 138 sediments include the Sykes Mountain Formation in the
Bighorn Basin, and the Fall River Formation in eastern
Powder River Basin. These transitional marine sediments are time equivalents of the lacustrine and fluvial Smoot
Formation adjacent to the fold-thrust belt. Figure 35 is a chart of the major tectonic and stratigraphie events in the overthrust belt and the foreland basin during the latest
Jurassic and Early Cretaceous.
As mentioned earlier, many authors have discussed the partitioning of the Early Cretaceous foreland basin (e.g.,
DeCelles, 1984; 1986; Kvale and Beck, 1985). The partitioning of the foreland basin in northern Wyoming and southern Montana is evidenced by the occurrence of dark, chert-bearing conglomerates at the base of the Cloverly and
Lakota formations in the Bighorn and Powder River basins, respectively, which may be traceable to the Ephraim
Conglomerate with similar dark chert, although the Ephraim
Conglomerate seems to be restricted only to the fold-thrust belt area west of the Prospect Thrust (Furer, 1970; Sippel,
1982). According to DeCelles (1984; 1986), during the earliest Cretaceous the basal conglomerate of the Kootenai
Formation over the Beartooth Uplift was reworked, transported eastward, and redeposited, forming the Pryor
Conglomerate of Moberly (1960). Thus the Beartooth uplift TECTONIC central western hoback basin some equivalent EVENTS ON southwesternmost idaho- wyoming,including (northern end of units else GEOLOûlC THE WE ST wyoming, THE north overthrust belt green river basin) where includ IN and the adjacent TIME over- central including fossil and western side of 'ng Farther thrust utah basin green river overthrust belt to the west in east in belt basin wyoming Wyoming LATE C«noma FRONT1ER FMy: CRET. nian -lOO MoWRY SH. ASPEN FM. Albion 2 ^ bakoVâ-" M smiths Fni-tpi^ a o
Aptian OC M u CO •Lplilta^îj•••••» »?%• • vo
N«o— cotnion W Por»- Initi tn londian •nov* mailt Lower Ephraim Fm. (sh.) on Paris and Wilard ' 5 lOxfor- Thrusts STUMP Redwaler Idian FM. Sh. Mbr.
Markedly Conglomeratic CFIuvlatile) I I Marin* and/or Brackish Sandstones and Finer coal Hiatus M elastics(piuvidtile) E3 Dominantly Lacustrine 1,2,3 Renamed movement on Paris Thrust
FIGURE 35. Sedimentation on the overthrust belt^and the adjacent foreland basin (modified from Wiltschko and Dorr, 1983) 140 represented one of the intrabasin source areas (Kvale,
1986). A volcanic highland in the Yellowstone/ northern
Absaroka region represents the source area of volcaniclastic sediments in the Lower Himes Member of the Cloverly
Formation, as indicated by paleocurrent data and abundant sand-size volcanic rock fragments, zoned plagioclase, and andesite clasts in these sediments (Kvale, 1986). Kvale
(1986) reported that the large, type I channel systems of the Greybull Sandstone were restricted to the regions north of the area of maximum uplift along the west flank of the
Bighorn Mountains. He added that the Greybull channels spread out from these regions, flowing west to southwest, and were preserved as isolated channel sandstone bodies separated from each other by several kilometers of thin, overbank deposits. Furthermore, he attributed the areal distribution of the Greybull channels to the effect of the structures associated with the Bighorn Mountains on sedimentation during the Early Cretaceous time.
Summary
The Wyoming Province was affected by the Sevier and
Laramide orogenic movements. The Sevier Orogeny began during Late Jurassic in the overthrust belt to the west.
The Laramide Orogeny began during Late Cretaceous in the 141 foreland on the east. After the middle Late Cretaceous, the two orogenies were overlapping in the time of movement.
Sedimentation within the early Sevier foreland basin was represented by the deposition of the different formations of the Gannett Group (fluvial and lacustrine) adjacent to the fold-thrust belt, and the fluvial Cloverly and Lakota formations in the Bighorn Basin and eastern Wyoming, respectively. The Early Cretaceous seaway transgressed from the Arctic Ocean southward at the beginning of Albian time due to further subsidence of the foreland. The transitional marine sediments include the Sykes Mountain and Fall River formations in the Bighorn and Powder River basins, respectively, while the lacustrine and fluvial Smoot
Formation was deposited adjacent to the fold-thrust belt.
The source areas that supplied the detritus to the foreland include the western highland of the overthrust belt and the eastern highlands of the Sioux Arch. The foreland basin has been tectonically partitioned by the Laramide Orogeny, and the intrabasin detritus was derived from an incipient
Beartooth uplift as well as a volcanic highland in the
Yellowstone/northern Absaroka region. 142
GENERAL SUMMARY AND CONCLUSIONS
1. The Sykes Mountain Formation consists of rusty interbedded and interlaminated sandstones and siltstones, and dark gray shales. It disconformably overlies the fluviatile Cloverly Formation and conformably underlies and interfingers with the marine Thermopolis Shale. Sandstones dominate the lower part while mudstones increase upward, forming an overall fining-upward sequence. The Sykes
Mountain Formation of the Bighorn Basin is correlated with the First Cat Creek Sand of central Montana, the Fall River
Formation of the Black Hills, the Plainview Sandstone Member of the Colorado Front Range, the Lower Naturita Formation of the Colorado Plateau, and the Smoot Formation of western
Wyoming.
2. Two sedimentary facies associations are recognized within the formation. They are the Quartz Arenite Facies
Association (QA Facies Association) and the Heterolithic
Facies Association (SM Facies Association). The two facies associations are arranged in the form of a fining-upward sequence in which the QA Facies Association grades upward to the SM Facies Association. The QA Facies Association is divided into seven facies according to differences in sedimentary structures. Sediments of the Sykes Mountain
Formation represent shallow subtidal, intertidal, and shallow marine environments. 143
3. Paleocurrent flow directions were almost in all directions, with the dominance of easterly and westerly directions, indicating that the sediments of the Sykes
Mountain Formation were derived mainly from the east and northeast as well as from the west.
4. Heavy minerals of the Sykes Mountain Formation indicate that its sediments were derived from metamorphic and sedimentary terrains which are likely the metamorphic rocks of the Canadian Shield and the Paleozoic sedimentary rocks of the Cordilleran region to the west.
5. Sedimentation during the Early Cretaceous in
Wyoming was mainly controlled by different episodes of movement along the Paris Thrust to the west as well as the
Laramide Orogeny within the foreland basin. 144
BIBLIOGRAPHY
Allen, J. R. L. 1982. Sedimentary structures; Their character and physical basis (Vol. II): Developments in sedimentology 30B. Elsevier, New York. 663p.
Allmendinger, R. W. and T. E. Jordan. 1981. Mesozoic evolution, hinterland of the Sevier orogenic belt. Geology 9:308-313.
Armstrong, R. L. 1968. Sevier orogenic belt in Nevada and Utah. Am. Assoc. Pet. Geol. Bull. 79(4):429-458.
Armstrong, F. C. and S. S. Oriel. 1965. Tectonic development of the Idaho-Wyoming thrust belt. Am. Assoc. Pet. Geol. Bull. 49:1847-1866.
Baird, G. C. 1978. Pebbly phosphorites in shale: A key to recognition of a widespread submarine discontinuity in the Middle Devonian of New York. J. Sed. Pet. 48; 545-555.
Banks, N. L. 1973. Tide-dominated offshore sedimentation. Lower Cambrian, North Norway. Sedimentology 20:213-228.
Bartsch-Winkler, S. and A. T. Ovenshine. 1984. Macrotidal subarctic environment of Turnagain and Kink Arms, Upper Cook Inlet, Alaska: sedimentology of the intertidal zone. J. Sed. Pet. 54:1221-1238.
Bartsch-Winkler, S. and H. R. Schmoll. 1984. Bedding types in Holocene tidal channel sequences. Kink Arm, Upper Cook Inlet, Alaska. J. Sed. Pet. 54(4):1239-1250.
Beard, D.C. and P. K. Weyl. 1973. Influence of texture on porosity and permeability of unconsolidated sand. Am. Assoc. Pet. Geol. Bull. 57:349-369.
Beaumont, C. 1978. The evolution of sedimentary basins on a viscoelastic lithosphere: Theory and examples. Royal Astronomical Society Geophysical Journal 55:471-497.
Beaumont, C. 1981. Foreland basins. Royal Astronomical Society Geophysical Journal 65:291-329. 145
Beck, R. A. 1985. Syntectonic sedimentation adjacent to Laramide basement thrusts, Rocky Mountain Foreland; Timing of deformation. Unpublished M. S. thesis. Iowa State University, Ames, Iowa.
Beck, R. A., C. F. Vondra, J. E. Filkins, and J. D. Olander. 1988. Syntectonic sedimentation and Laramide basement thrusting. Rocky Mountain Foreland: Timing and deformation. %n C. J. Schmidt and W. J. Perry (eds.), Interaction of the Rocky Mountain Foreland and Cordilleran thrust belt. Geol. Soc. Am. Memoir (in press).
Berg, R. R. 1962. Mountain flank thrusting in Rocky Mountain foreland, Wyoming and Colorado. Am. Assoc. Pet. Geol. Bull. 46:2019-2032.
Berg, R. R. 1981. Review of thrusting in the Wyoming foreland. Contrib. to Geol., Univ. of Wyo. 19(2);93-104.
Berg, R. R. 1983. Geometry of the Wind River Thrust, Wyoming, p. 257-262. _In J. D. Lowell (ed.). Rocky Mountain Foreland Basins and Uplifts, Rocky Mountains Association of Geologists, Denver, Colorado.
Bhatia, M. R. 1985. Composition and classification of Paleozoic flysch mudrocks of eastern Australia: Implications in provenance and tectonic setting interpretation. Sedimentary Geology 41:249-268.
Biscaye, P. E. 1965. Mineralogy and sedimentation of recent deep-sea clay in the Atlantic Ocean and adjacent seas and oceans. Geol. Soc. Am. Bull. 76:803-832.
Blackstone, D. L. 1981. Comparison as an agent in deformation of the east-central flank of the Bighorn Mountains, Sheridan and Johnson Counties, Wyoming. Contributions to Geology (University of Wyoming) 19(2):105-122.
Blackstone, D. L., Jr. 1983. Laramide compressional tectonics, southeastern Wyoming. Contributions to Geology (University of Wyoming) 22(l):l-38.
Blakey, R. C. 1984. Marine sand-wave complex in the Permian of central Arizona. J. Sed. Pet. 54:29-51.
Blatt, H. 1979. Diagenetic processes in sandstones. SEPM Special Publication 26:141-157. 146
Blatt, H. 1982. Sedimentary Petrology. W. H. Freeman and Company, San Francisco. 564 p.
Blatt, H., G. Middleton and R. Murray. 1972. Origin of sedimentary rocks. Prentice-Hall, Inc., Englewood Cliffs, New Jersey.
Boles, J. R. and S. G. Franks. 1979. Clay diagenesis in Wilcox sandstones of southwest Texas; Implications of smectite diagenesis on sandstone cementation. J. Sed. Pet. 49:55-70.
Bouma, A. H., M. L. Rappeport, R. C. Orlando, and M. A. Hampton. 1980. Identification of bedforms in Lower Cook Inlet, Alaska. Sedimentary Geology 26:157-177.
Bown, T. M. 1980. Summary of latest Cretaceous and Cenozoic sedimentary, tectonic, and erosional events. Bighorn Basin, Wyoming, p. 25-32. %n P. D. Ginerich (ed.). Early Cenozoic paleontology and stratigraphy of the Bighorn Basin, Wyoming. University of Michigan Papers on Paleontology No. 24.
Brenchley, P. J. 1985. Storm influenced sandstone beds. Modern Geology 9:369-396.
Brett, G. W. 1955. Cross-bedding in the Baraboo Quartzite of Wisconsin. J. Geology 63:143-148.
Briggs, G. and L. M. Cline. 1967. Paleocurrents and source areas of Late Paleozoic sediments of the Ouachita Mountains. J. Sed. Pet. 37:985-1000.
Brookfield, M. E. 1977. The origin of bounding surfaces in ancient aeolian sandstones. Sedimentology 24:303-332.
Burk, C. A. 1956. Subsurface stratigraphy of the Pre- Niobrara formations in the Wind River Basin, Central Wyoming. P. 23-33. _In Wyo. Geol. Assoc. Guidebook 1. Wyo. Geol. Assoc., Casper, Wyoming.
Burk, C. A. 1957. Stratigraphie summary of the non-marine Upper Jurassic and Lower Cretaceous of Wyoming. Wyo. Geol. Assoc. Guidebook 12:55-62.
Cazeau, C. J, 1960. Cross-bedding directions in Upper Triassic sandstones of west Texas. J. Sed. Pet. 30:459-465. 147
Chamberlain, R. T. 1945. Basement control in Rocky Mountain deformation. Am. J. Sci. 243-A:98-110.
Chisholm, W. A. 1963. The Petrology of Upper Jurassic and Lower Cretaceous strata of the Western Interior. P. 71-86. _In Wyo. Geol. Assoc. and Billings Geol. Soc. Guidebook to northern Powder River Basin. Laramie, Wyoming.
Chisholm, W. A. 1970. Depositional history and petrography of two Fall River Formation sandstone reservoirs in the eastern Powder River Basin and their relationship to oil entrapment. Wyo. Geol. Assoc. Guidebook 22:169-185.
Clifton, H. E. 1983. Discrimination between subtidal and intertidal facies in Pleistocene deposits, Willapa Bay, Washington. J. Sed. Pet. 53(2);353-369.
Clifton, H. E., R. E. Hunter, and R. L. Phillips. 1971. Depositional structures and processes in the nonbared high energy nearshore. J. Sed. Pet. 41:651-670.
Collinson, J. D. 1970. Bedforms of the Tana River, Norway. Geogr. Annaler 52-A;31-56.
Collinson, J. D. and D. P. Thompson. 1982. Sedimentary structures. George Allen & Unwin, London.
Comstock, T. B. 1873. On the geology of the northwestern Wyoming Expedition 1873. Am. J. Sci. and Arts 3rd ser G;426-432.
Coney, P. J. 1972. Cordilleran tectonics and North American plate motion. Am. J. Sci. 272:603-628.
Craft, J. H. and J. S. Bridge. 1987. Shallow-marine sedimentary processes in the Late Devonian Catskill Sea, New York State. Geol. Soc. Am. Bull. 98:338-355.
Darton, N. H. 1904. Comparison of the stratigraphy of the Black Hills, Bighorn Mountains and Rocky Mountain front range. Geol. Soc. Am. Bull. 15:379-448.
Darton, N. H. 1906a. Geology of the Bighorn Mountains. U. S. Geol. Surv. Prof. Paper 51.
Darton, N. H. 1906b. Cloud Peak-Fort McKinney, Wyoming. U. S. Geol. Surv. Geol. Atlas. Folio 142. 148
DeCelles, P. G. 1984. Sedimentation and disgenesis in a tectonically partitioned, nonmarine foreland basin; The Lower Cretaceous Kootenai Formation, southwestern Montana. Unpublished Ph.D. dissertation. Indiana University, Bloomington, Indiana.
DeCelles, P. G. 1986. Sedimentation in a tectonically partitioned, nonmarine foreland basin: The Lower Cretaceous Kootenai Formation, southwestern Montana. Geol. Soc. Am. Bull. 97(8);911-931.
Dickinson, W. R. 1974. Plate tectonics and sedimentation. SEPM Special Publication 22:1-27.
Dickinson, W. R., M. A. Klute, M. J. Hayes, S. U. Janecke, E. R. Lundin, M. A. McKittrick, and M. D. Olivares. 1988. Paleogeographic and paleotectonic setting of Laramide sedimentary basins in the central Rocky Mountain region. Geol. Soc. Am. Bull. 100(7):1023-1039.
Doe, T. W. and R. H. Dott, JR. 1980. Genetic significance of deformed cross-bedding with examples from the Navajo and Weber Sandstones of Utah. J. Sed. Pet. 50:793-812.
Dorr, J. A., Jr., D. R. Spearing, and J. R. Steidtmann. 1977. Deformation and deposition between a foreland uplift and an impinging thrust belt. Geol. Soc. Am. Special Paper 177:1-82.
Dott, R. H., Jr. 1973. Paleocurrent analysis of trough cross stratification. J. Sed. Pet. 43:779-783.
Dott, R. H., Jr. and J. Bourgeois. 1982. Hummocky stratification; Significance of its variable bedding sequences. Geol. Soc. Am. Bull. 93:663-680.
Dunn, D. J. 1986. The geology of the Bear Creek area, Wyoming. Unpublished M.S. thesis. Iowa State University, Ames, Iowa.
Durkee, S. K. 1980. Depositional environments of the Lower Cretaceous Smiths Formation within a portion of the Idaho-Wyoming thrust belt. Wyo. Geol. Assoc. Guidebook 31:101-116.
Eicher, D. L. 1958. The Thermopolis Shale in Eastern Wyoming. Wyo. Geol. Assoc. Guidebook 13:79-83. 149
Eicher, D. L. 1960. Stratigraphy and micropaleontology of the Thermopolis Shale. Yale Univ., Peabody Mus. Nat. Hist. Bull. 15:126.
Eicher, D. L. 1962. Biostratigraphy of the Thermopolis, Muddy, and Shell Creek Formations. Wyo. Geol. Assoc. Guidebook 17:72-93.
Ekdale, A. A., R. G. Bromley and S. G. Pemberton. 1984. Ichnology: Trace fossils in sedimentology and stratigraphy. SEPM Short Course No. 15. SEPM, Tulsa, Oklahoma.
Eldridge, G. H. 1894. A geological reconnaissance of northwest Wyoming. U. S. Geol. Surv. Bull. 119:1-69.
Evans, G. 1965. Intertidal flat sediments and their environments of deposition in The Wash. Q. Geol. Soc. London 121:209-245.
Evans, G. 1975. Intertidal flat deposits of the Wash, western margin of the North Sea. P. 13-20. _In R. N. Ginsburg (ed.). Tidal deposits. Springer-Verlag, New York, N. Y.
Eyer, J. A. 1969. Gannett Group of western Wyoming and southeastern Idaho. Am. Assoc. Pet. Geol. Bull. 53(7);1368-1390.
Fanshawe, J. R. 1971. Structural evolution of the Bighorn Basin. Wyo. Geol. Assoc. Guidebook 23:35-37.
Faulkner, G. L. 1956. Subsurface stratigraphy of the Pre- Niobrara formations along the western margin of the Powder River Basin, Wyoming. P. 35-42. In Wyo. Geol. Assoc. Guidebook 1: Wyoming Stratigraphy. Wyo. Geol. Assoc., Casper, Wyoming.
Fisher, C. A. 1906. Geology and water resources of the Bighorn Basin, Wyoming. U. S. Geol. Surv. Prof. Paper 53:1-72.
Fothergill, C. A. 1955. The cementation of oil reservoir sands and its origin. Fourth World Petroleum Congress Sec. 1/B, Paper 1:301-314.
Franks, P. C., G. L. Coleman, N. Plummer, and W. K. Hamplin 1959. Cross-stratification, Dakota Sandstone (Cretaceous), Ottawa County, Kansas. Geol. Surv. Kansas Bull. 134:223-238. 150
Purer, L. C. 1970. Petrology and stratigraphy of nonmarine Upper Jurassic-Lower Cretaceous rocks of western Wyoming and southeastern Idaho. Am. Assoc. Pet. Geol. Bull. 54:2282-2302.
Gott, G. B., D. E. Wolcott, C. G. Bowles. 1974. Stratigraphy of the Inyan Kara Group and localization of uranium deposits, southern Black Hills, South Dakota and Wyoming. U. S. Geol. Surv. Prof. Paper 763.
Gries, J. P. 1962. Lower Cretaceous stratigraphy of South Dakota and the eastern edge of the Powder River Basin. Wyo. Geol. Assoc. Guidebook 17:163-172.
Gries, R. 1983. North-south compression of Rocky Mountain foreland structures, p. 9-32. %n J. D. Lowell and R. Gries (eds.). Rocky Mountain Foreland Basins and Uplifts. Rocky Mtn. Assoc. Geol. Symp., 1983.
Groot, J. J. 1955. Sedimentary petrology of the Cretaceous sediments of northern Delaware in relation to paleogeographic problems. Delaware Geol. Surv. Bull. 5:1-157.
Grose, L. T. 1972. Geologic atlas of the Rocky Mountain Region, United States of America. U.S.A. Hirschfield Press, Denver, Colorado.
Hamilton, R. D. 1973. An investigation of two aspects of phosphorite genesis. Unpublished M. S. thesis. Iowa State University, Ames, Iowa.
Hares, C. J. 1917. Gastroliths in the Cloverly Formation. J. Wash. Acad. Sci. 7:429.
Harms, J. C., J. B. Southard, D. R. Spearing, and R. G. Walker. 1975. Depositional environments as interpreted from primary sedimentary structures and stratification sequences. SEPM short course 2. 161 p. SEPM, Tulsa, Oklahoma.
Harms, J. C., J. B. Southard, and R. G. Walker. 1982. Structures and sequences in clastic rocks. SEPM short course 9. 249 p. SEPM, Tulsa, Oklahoma.
Haun, J, D. 1959. Lower Cretaceous stratigraphy of Colorado. Wyo. Geol. Assoc. Guidebook 17:1-8. 151
Haun, J. D. 1963. Stratigraphy of Dakota Group and relationship to petroleum occurrence, northern Denver Basin. Rocky Mountain Assoc. Geol. Guidebook 14:119-134.
Haun, J. D. and J. A. Barlow, Jr. 1962. Lower Cretaceous stratigraphy of Wyoming. Wyo. Geol. Assoc. Guidebook 17:15-22.
Haun, J. D. and R. J. Weimer. 1960. Cretaceous stratigraphy of Colorado. P. 58-65. In Guide to the Geology of Colorado: Geol. Soc. Am., Rocky Mtn. Assoc. Geol., and Colorado Sci. Soc.
Heald, M. T. 1955. Stylolites in sandstones. J. Geology 63:101-114.
Hewett, D. F. 1914. The Shoshone River Section, Wyoming. U. S. Geol. Surv. Bull. 541:89-112.
Hewett, D. F., and C. T. Lupton. 1917. Anticlines in the southern part of the Bighorn Basin, Wyoming. U. S. Geol. Surv. Bull. 656:3-192.
Hiltabrand, R. R., R. E. Ferrell, and G. K. Billings. 1973. Experimental diagenesis of Gulf Coast argillaceous sediment. Am. Assoc. Pet. Geol. Bull. 57:338-348.
Hintze, F. F., Jr. 1915. The Basin and Greybull oil and gas fields. Wyo. Geol. Surv. Bull. 10:1-62.
Holm, M. R., W. C. James and L. J. Suttner. 1977. Comparison of the Peterson and Draney Limestones, Idaho and Wyoming, and the calcareous members of the Kootenai Formation, western Montana. Wyo. Geol. Assoc. Guidebook 29:259-270.
Hower, J., E. V. Eslinger, M. E. Hower, and E. A. Perry. 1976. Mechanism of burial metamorphism of argillaceous sediments: Mineralogical and chemical evidence. Geol. Soc. Am. Bull. 87:725-737.
Hunter, R. E. and H. E. Clifton. 1982. Cyclic deposits and hummocky cross-stratification of probable storm origin in Upper Cretaceous rocks of the Cape Sebastian area, Southwestern Oregon. J. Sed. Pet. 52:127-143.
Johnson, H. D. 1978. Shallow siliciclastic seas. P. 207-258. In H. G. Reading (ed.). Sedimentary environments and facies. Elsevier, New York. 152
Jordan, T. E. 1981. Thrust loads and foreland basin evolution. Am. Assoc. Pet. Geol. Bull. 65(12);2506-2520.
Kauffman, E. G. 1977. Geological and biological overview: Western Interior Cretaceous basin. The Mountain Geologist 14:75-99.
Kent, H. C. 1972. Review of phanerozoic history. P. 57-59. %n Rocky Mountain Association of Geologists. Geologic Atlas of the Rocky Mountain Region. Hirschfield Press, Denver, Colorado.
King, D. T., Jr. and H. S. Chafetz. 1983. Tidal-flat to shallow-shelf deposits in the Cap Mountain Limestone Member of the Riley Formation, Upper Cambrian of central Texas. J. Sed. Pet. 53Q):261-273.
Klein, G. D. 1963. Bay of Fundy intertidal zone sediments. J. Sed. Pet. 33:844-854.
Klein, G. D. 1967. Paleocurrent analysis in relation to modern marine sediment dispersal patterns. Am. Assoc. Pet. Geol. Bull. 51:366-382.
Klein, G. D. 1970a. Depositional and dispersal dynamics of intertidal sand bars. J. Sed. Pet. 40:1095-1127.
Klein, G. D. 1970b. Tidal origin of a Pre Cambrian Quartzite-the Lower Fine Grained Quartzite (Middle Dalradian) of Islay, Scotland. J. Sed. Pet. 40:973-985.
Klein, G. D. 1971. A sedimentary model for determining paleotidal range. Geol. Soc. Am. Bull. 82:2585-2592.
Klein, G. D. 1972. Determination of paleotidal range in clastic sedimentary rocks. Int. Geol. Congr. 24(6):397-405.
Klein, G. D. 1975. Paleotidal range sequences. Middle Member, Wood Canyon Formation (Late Precambrian), eastern California and western Nevada. P. 1171-1177. In R. N. Ginsburg (ed.). Tidal deposits. Springer- Verlag, New York, N. Y.
Knappan, R. S. and G. F. Moulton. 1930. Geology and mineral resources of parts of Carbon, Bighorn, Yellowstone, and Stillwater Counties, Montana. U. S. Geol. Surv. Bull. 822A:l-70. 153
Kozimko, L. M. 1977. Geology of the Greybull North Quadrangle, Wyoming. Unpublished M. S. thesis. Iowa State University, Ames, Iowa.
Kreisa, R. D. 1981. Storm-generated sedimentary structures in subtidal marine facies with examples from the Middle and Upper Ordovician of southwestern Virginia. J. Sed. Pet. 51:823-848.
Kvale, E. P. 1986. Paleoenvironments and tectonic significance of the Upper Jurassic Morrison and Lower Cretaceous Cloverly formations. Bighorn Basin, Wyoming. Unpublished Ph.D. dissertation. Iowa State University, Ames, Iowa.
Kvale, E. P. and R. A. Beck. 1985. Thrust controlled sedimentation patterns of the earliest Cretaceous nonmarine sequence of the Montana-Idaho-Wyoming Sevier foreland basin. Geol. Soc. Am. Abstr. Prog. 17:636.
Ladd, R. E. 1979. The geology of Sheep Canyon Quadrangle, Wyoming. Unpublished M.S. thesis. Iowa State University, Ames, Iowa.
Lee, W. T. 1923. Continuity of some oil-bearing sands of Colorado and Wyoming. U. S. Geol. Surv. Bull. 751-A:l-22.
Lee, W. T. 1927. Correlation of geologic formations between east-central Colorado, central Wyoming and southern Montana. U. S. Geol. Surv. Prof. Paper 149:1-57.
Levandowski, D. W., M. E. Kaley, S. R. Silverman, and R. G. Smalley. 1973. Cementation in Lyons Sandstone and its role in oil accumulation, Denver Basin, Colorado. Am. Assoc. Pet. Geol. Bull. 57:2217-2244.
Love, J. D. 1960. Cenozoic sedimentation and crustal movement in Wyoming. Am. J. Sci. 258-A:204-214.
Lupton, C. T. 1916. Oil and gas near Basin, Bighorn County, Wyoming. U. S. Geol. Surv. Bull. 621L:156-190.
Ly, C. K. 1982. Sedimentology of nearshore marine bar sequences from a Paleozoic depositional regressive shoreline deposit of the central coast of Ghana, west Africa. J. Sed. Pet. 52:199-208. 154
MacKenzie, D. B. and D. M. Poole. 1962. Provenance of Dakota Group sandstones of the Western Interior. Wyo. Geol. Assoc. Guidebook 17:62-71.
MacKenzie, D. E. 1968. Studies for students; Sedimentary features of Alameda Avenue Cut, Denver, Colorado. Mtn. Geologist 5(1):3-13.
MacKenzie, F. T. and R. Gees. 1971. Quartz: synthesis at earth-surface conditions. Science 173:533-535.
MacKenzie, F. T. and J. D. Ryan. 1962. Cloverly-Lakota and Fall River paleocurrents in the Wyoming Rockies. Wyo. Geol. Assoc. Guidebook 17:44-61.
Manahl, K. A. 1981. Geology of the Shell Quadrangle, Wyoming. Unpublished M.S. thesis. Iowa State University, Ames, Iowa.
Mantzios, C. 1986. The paleosols of the Upper Jurassic Morrison and Lower Cretaceous Cloverly formations. Northeast Bighorn Basin, Wyoming. Unpublished M.S. thesis. Iowa State University, Ames, Iowa.
McCave, I. N. 1970. Deposition of fine-grained suspended sediment from tidal currents. J. Geophys. Res. 75:4151-4159.
McConnell, D. 1935. Spherulitic concretions of dahllite from Ishawooa, Wyoming. Am. Mineralogist 20:693-698.
McConnell, D. 1960. The crystal chemistry of dahllite. Am. Mineralogist 45:209-216.
Meek, F. B. and F. V. Hayden. 1856a. Descriptions of new species of Acephala and Gastropoda, from the Tertiary formations of Nebraska Territory, with some general remarks on the geology of the country about the sources of the Missouri River. Proc. Acad. Nat. Sci. Phila. 8:111-126.
Meek, F. B. and F. V. Hayden. 1856b. Descriptions of new fossil species of mullusca collected by Dr. F. V. Hayden, in Nebraska Territory; together with a complete catalogue of all the remains of invertebrata hitherto described and identified from the Cretaceous and Tertiary formations of that region. Proc. Acad. Nat. Sci. Phila. 8:265-286. 155
Meek, F. B. and F. V. Hayden. 1858. Descriptions of new organic remains collected in Nebraska Territory in the year 1857, by Dr. F. V. Hayden, geologist to the exploring expedition under the command of Lieut. G. K. Warren, Top. Engr. U. S. Army, together with some remarks on the geology of the Black Hills and portions of the surrounding country. Proc. Acad. Nat. Sci. Phila. 10:41-59.
Meek, F. B. and F. V. Hayden. 1861. Descriptions of new Lower Silurian (Primordial), Jurassic, Cretaceous, and Tertiary fossils, collected in Nebraska, by the exploring expedition under the command of Capt. Wm. F. Reynolds, U. S. Top. Eng.; with some remarks on the rocks from which they were obtained. Proc. Acad. Nat. Sci. Phila. 13:415-447.
Miall, A. D. 1977. A review of the braided river depositional environment. Earth Sci. Rev. 13:1-62.
Miall, A. D. 1978. Lithofacies types and vertical profile models in braided river deposits: A summary. _In A. D. Miall (ed.). Fluvial Sedimentology. Mem. Can. Soc. Petrol. Geol. 5:597-604. Calgary, Alberta, Canada.
Michelson, P. C. and R. H. Dott, Jr. 1973. Orientation analysis of trough cross-stratification in Upper Cambrian sandstones of western Wisconsin. J. Sed. Pet. 43:784-794.
Miller, M. F. 1984. Bioturbation of intertidal quartz-rich sands: a modern example and its sedimentologic and paleoecologic implications. J. Geology 92:201-216.
Mills, N. K. 1956. Subsurface stratigraphy of the Pre- Niobrara formations in the Bighorn Basin, Wyoming. P. 9-22. _In Wyo. Geol. Assoc. Guidebook 1. Wyoming Stratigraphy, Casper, Wyoming.
Mirsky, A. 1962a. Stratigraphie sections of Upper Jurassic and Lower Cretaceous rocks in the southern Bighorn Mountains, Wyoming. Wyoming Geol. Surv. Rept. Invest. 8:1-33.
Mirsky, A. 1962b. Stratigraphy of nonmarine Upper Jurassic and Lower Cretaceous rocks, southern Bighorn Mountains, Wyoming. Am. Assoc. Pet. Geol. Bull. 46:1653-1680. 156
Moberly, R., Jr. 1956. Mesozoic Morrison, Cloverly, and Crooked Creek Formations, Bighorn Basin, Wyoming and Montana. Unpublished Ph.D. dissertation. Princeton University.
Moberly, R., Jr. 1960. Morrison, Cloverly and Sykes Mountain Formations, northern Bighorn Basin, Wyoming and Montana. Geol. Soc. Am. Bull. 71:1137-1176.
Moberly, R., Jr. 1962. Lower Cretaceous history of the Bighorn Basin, Wyoming. Wyo. Geol. Assoc. Guidebook 17:94-101.
Moore, D. G. and P. C. Scruton. 1957. Minor internal structures of some recent unconsolidated sediments. Am. Assoc. Pet. Geol. 41:2723-2751.
Mowbray, T. D. and M. J. Visser. 1984. Reactivation surfaces in subtidal channel deposits, Oosterschelde, southwest Netherlands. J. Sed. Pet. 54:811-824.
Noggle, K. S. 1986. Stratigraphy and structure of the Leavitt Reservoir Quadrangle, Bighorn County, Wyoming. Unpublished M.S. thesis. Iowa State University, Ames, Iowa.
Oriel, S. S. and F. C. Armstrong. 1966. Times of thrusting in Idaho-Wyoming thrust belts: Reply. Am. Assoc. Pet. Geol. 50:2614-2621.
Oriel, S. S. and L. B. Piatt. 1979. Younger-over-older thrust plates in southeastern Idaho. Geol. Soc. Am. Abs. with Programs 11:298. (abstr.).
Osgood, R. G., Jr. 1970. Trace fossils of the Cincinnati area. Palaeontographica Am. 6 (41):281-444.
Ostrom, J. H. 1970. Cloverly Formation, stratigraphy and paleontology. Yale Univ. Peabody Museum of Natural Hist. Bull. 35:1-234.
Pasierbiewicz, K. W. 1982. Experimental study of cross- strata development on an undulatory surface and implications relative to the origin of flaser and wavy bedding. J. Sed. Pet. 52:769-778.
Pettijohn, F. J, 1957. Paleocurrents of Lake Superior Precambrian quartzites. Geol. Soc. Am. Bull. 68:469-480. 157
Pettijohn, F. J. 1962. Paleocurrents and paleogeography. Am. Assoc. Pet. Geol. Bull. 46:1468-1493.
Pierce, W. G. 1948. Geologic and structure contour map of the Basin-Greybull area, Bighorn County, Wyoming. U. S. Geol. Surv. Oil and Gas Invest. Map 77.
Pierce, W. G., and D. A. Andrews. 1940. Geology of oil and coal resources of the region south of Cody, Park County, Wyoming. U. S. Geol. Surv. Bull. 921B;99-180.
Pittman, E. D. 1972. Diagenesis of quartz in sandstones as revealed by scanning electron microscopy. J. Sed. Pet. 42:507-519.
Potter, P. E. and J. S. Olson. 1954. Variance components of cross-bedding direction in some basal Pennsylvanian sandstones of the Eastern Interior Basin: Geological application. J. Geology 62:50-73.
Potter, P. E. and F. J. Pettijohn. 1977. Paleocurrents and basin analysis. Springer-Verlag, New York. 425 p.
Potter, P. E. and W. A. Pryor. 1961. Dispersal centers of Paleozoic and later clastics of the Upper Mississippi Valley and adjacent areas. Geol. Soc. Am. Bull. 72:1195-1250.
Potter, P. E. and R. Siever. 1956. Sources of basal Pennsylvanian sediments in the Eastern Interior Basin: 1. Crossbedding. J. Geology 64:225-244.
Powers, M. C. 1953. A new roundness scale for sedimentary particles. J. Sed. Pet. 23:117-119.
Prucha, J. J., J. A. Graham, and R. P. Nickelson. 1965. Basement-controlled deformation in Wyoming Province of Rocky Mountains Foreland. Am. Assoc. Pet. Geol. Bull. 49(7):966-992.
Raaf, J. F. M. De and J. R. Boersma. 1971. Tidal deposits and their sedimentary structures. Geol. Mijnb. 50(3):479-504.
Ramsay, J. G. 1961. The effects of folding upon the orientation of sedimentation structures. J. Geology 69:84-100.
Reading, H. G. 1978. Sedimentary environments and facies. Blackwell Scientific Publications, Oxford. 569p. 158
Reeside, J. B., Jr. 1923. Fauna of the so-called Dakota Formation of northern central Colorado and its equivalent in southeastern Wyoming. U. S. Geol. Surv. Prof. Paper 131-H;199-207.
Reineck, H. E. 1960a. Uber die Entstehung von linsen und flaserschichten. Abhandl. Deut. Akad. Wiss. 1:370-374.
Reineck, H. E. 1960b, Uber zeitlucken in rezenten flachsee-sedimenten. Geol. Rundschau. 49:149-161.
Reineck, H. E. 1963. Sedimentgefuge im Bereich der Sudliche Nordsee. Abh. Senchenb. Naturforch. Ges. 505:1-138.
Reineck, H. E. 1967a. Parameter von schichtung und bioturbation. Geologische Rundschau. 56:420-438.
Reineck, H. E. 1967b. Layered sediments of tidal flats, beaches, and shelf bottoms of the North Sea. In G. H. Lauff, (ed.), Estuaries. Am. Assoc. Advanc. Sci. 83:191-206. Washington.
Reineck, H. E. 1975. German North Sea tidal flats. P. 5-12. %n R. N. Ginsburg (ed.), Tidal deposits: A casebook of recent examples and fossil counterparts. Springer-Verlag, New York.
Reineck, H. E. and I. B. Singh. 1980. Depositional sedimentary environments (2nd ed.). Springer-Verlag, New York.
Reineck, H. E. and F. Wunderlich. 1968. Classification and origin of flaser and lenticular bedding. Sedimentology 11:99-104.
Reppe, C. C. 1981. The geology of Devil's Kitchen Quadrangle, Wyoming. Unpublished M.S. thesis. Iowa State University, Ames, Iowa.
Rhoads, D. C. 1967. Biogenic reworking of intertidal and subtidal sediments in Barnstable Harbor and Buzzards Bay, Massachusetts. J. Geology 75:461-476.
Rhoads, D. C. 1975. The paleoecological and environmental significance of trace fossils. P. 147-160. %n R. W. Frey (ed.). The study of trace fossils. Springer Verlag, New York. 159
Richards, P. W. 1955. Geology of the Bighorn Canyon-Hardin area, Montana and Wyoming. U. S. Geol. Surv. Bull. 1026:1-96.
Rittenhouse, G. 1943. A visual method of estimating two- dimensional sphericity. J. Sed. Pet. 13:79-81.
Rogers, A. F. 1912. Dahllite (podolite) from Tonapah, Nevada, voelckerite a new calcium phosphate; remarks on the chemical composition of apatites and phosphate rock. Am. J. Sci. 33:475-482.
St. John, 0. 1883. Report on the geology of the Wind River District. U. S. Geol. Geogr. Surv. Terr. Wyoming and Idaho for the year 1878. U. S. Gov. Printing Office, Washington, D.C.
Schedl, A. and D. V. Wiltschko. 1980. Isostatic effects of a moving thrust terraine. EOS 61:360.
Schmidt, V. and D. A. McDonald. 1979. The role of secondary porosity in the course of sandstone diagenesis. SEPM Special Publication 26:175-207.
Schwartz, R. K. 1983. Broken Early Cretaceous foreland basin in southwestern Montana: Sedimentation related to tectonism. p. 159-184. %n R. R. Powers (ed.). Geologic studies of the Cordilleran thrust belt. Rocky Mountain Association of Geologists. Denver, Colorado.
Seilacher, A. 1953. Studien zur palichnologie I, Uber die methoden der palichnologie. Neues Jahrb. palaont. Abh. 98:87-124.
Seilacher, A. 1960. Stromungsanzeichen im Hunsruckschief- er. Notizbl. hess. Landesamtes Bodenforsch. Wiesbaden 88:88-106.
Selley, R. C. 1982. An introduction to sedimentology. Academic Press, London.
Siever, R. 1957. The silica budget in the sedimentary cycle. Am. Mineralogist 42:821-841.
Siever, R. 1959. Petrology and geochemistry of silica cementation in some Pennsylvanian sandstones, p. 55-79. %n SEPM Spec. Publication 7: Silica in sediments. SEPM, Tulsa, Oklahoma. 160
Simpson, S. 1975. Classification of trace fossils. P. 39-54. %n R. W. Frey (éd.), The study of trace fossils. Springer-Verlag, Berlin.
Singer, A. 1980. The paleoclimatic interpretation of clay minerals in soils and weathering profiles. Earth Sci. Rev. 15:303-326.
Singer, A. 1984. The paleoclimatic interpretation of clay minerals in sediments-a review. Earth Sci. Rev. 21:251-293.
Singler, C. R. and M. D. Picard. 1979. Paleocurrent indicators and other sedimentary structures in the Brule Formation (Oligocene), northwest Nebraska. The Wyo. Geol. Assoc. Earth Science Bull. 12:17-25.
Sippel, K. N. 1982. Depositional and tectonic setting of the Ephraim Formation, southeastern Idaho and western Wyoming. Unpublished M.S. thesis. University of Wyoming, Laramie.
Sloss, L. L. and D. E. Feray. 1948. Microstylolites in sandstone. J. Sed. Pet. 18:3-13.
Soliman, H. E. and C. F. Vondra. 1988. Stratigraphy and environment of deposition of Lower Cretaceous Sykes Mountain, Bighorn Basin, Wyoming. Geol. Soc. Am. Abstr. Prog. 20:129.
Stanley, D. J. and D. J. P. Swift (eds.). 1976. Marine sediment transport and environmental management. Wiley, New York. 602 p.
Stearns, D. W. 1971. Mechanisms of drape folding in the Wyoming Province. Wyo. Geol. Assoc. Guidebook 23:125-143.
Stearns, D. W. 1975. Laramide basement deformation in the Bighorn Basin-The controlling factor for structures in the layered rocks. Wyo. Geol. Assoc. Guidebook 27:149-158.
Stearns, D. W. 1978. Faulting and forced folding in the Rocky Mountain foreland. P. 1-38. In V. Matthews (ed.), Laramide folding associated with basement block faulting in the western United States. Geol. Soc. Am. Memoir 151. 161
Suttner, L. J., R. K. Schwartz, and W. C. James. 1981. Late Mesozoic to Early Cenozoic foreland sedimentation in southwest Montana. P. 93-103. In Field Conf. southwest Montana. Montana Geol. Soc., Billings.
Swett, K., G. D. Klein, and D. E. Smit. 1971. A cambrian tidal sand body-the Eriboll Sandstone of northwest Scotland; an ancient-recent analog. J. Geology 79:400-415.
Swift, D. J. P., A. G. Figueiredo, JR., G. L. Freeland, and G. F. Oertel. 1983. Hummocky cross-stratification and megaripples; A geological double standard? J. Sed. Pet. 53(4);1295-1317.
Tankard, A. J. and D. k. Hobday. 1977. Tide-dominated back-barrier sedimentation. Early Ordovician Cape Basin, Cape Peninsula, South Africa. Sedimentary Geology 18:135-159.
Tanner, W. P. 1955. Paleogeographic reconstruction from cross-bedding studies. Am. Assoc. Pet. Geol. Bull. 39:2471-2483.
Tanner, W. F. 1958. An occurrence of flat-topped ripple marks. J. Sed. Pet. 28:95-96.
Taylor, J. m. 1950. Pore-space reduction in sandstones. Am. Assoc. Pet. Geol. Bull. 34:701-716.
Terwindt, J. H. J, 1971. Sand waves in the southern bight of the North Sea. Mar. Geology 10:51-67.
Terwindt, J. H. J. and H. N. C. Breusers. 1972. Experiments on the origin of flaser, lenticular and sand-mud alternating bedding. Sedimentology 19:85-98,
Todd, T. W. 1963. Post-depositional history of Tensleep Sandstone (Pennsylvanian), Bighorn Basin, Wyoming. Am. Assoc. Pet. Geol. Bull. 47:599-616.
Towe, K. M. 1962. Diagenesis of clay minerals as a possible source of silica cement in sedimentary rocks. J. Sed. Pet. 32:26-28.
Van Straaten, L. M. J. U. 1952. Biogene textures and the formation of shell beds in the Dutch Madden Sea. Koninkl. Ned. Akad. Wetenschap Proc., Ser. B 55:500-516. 162
Van Straaten, L. M. J. U. 1954a. Sedimentology of Recent tidal flat deposits and the Psammites du condroz (Devonian). Geol. Mijnbouw 16:25-47.
Van Straaten, L. M. J. U. 1954b. Composition and structure of Recent marine sediments in the Netherlands. Leidse Geologische Mededelingen. 19:1-110.
Van Straaten, L. M. J. U. 1961. Sedimentation in tidal flat areas; Alberta Soc. Petrol. Geol. J. 9:203-226.
Van Straaten, L. M. J. U. and Ph. H. Kuenen. 1958. Tidal action as a cause of clay accumulation. J. Sed. Pet. 28:406-413.
Von Englehardt, W. 1967. Interstitial solutions and diagenesis in sediments. P. 503-521. %n Developments in sedimentology 8: Diagenesis in sediments. Elsevier Pub. Co., Amsterdam.
Waage, K. M. 1955. Dakota Group in northern Front Range Foothills, Colorado. U. S. Geol. Surv. Prof. Paper 274-B:15-51.
Waage, K. M. 1958. Regional aspects of Inyan Kara stratigraphy. Wyo. Geol. Assoc. Guidebook 13:71-76.
Waage, K. M. 1959a. Stratigraphy of the Inyan Kara Group in the Black Hills. U. S. Geol. Surv. Bull. 1081-8:11-90.
Waage, K. M. 1959b. Dakota stratigraphy along the Colorado Front Range. Rocky Mountain Assoc. Geol. Guidebook 11:115-123.
Waage, K. M. 1961. Stratigraphy and rafrectory clayrocks of the Dakota Group along the northern Front Range, Colorado. U. S. Geol. Surv. Bull. 1102:154.
Waldschmidt, W. A. 1941. Cementing materials in sandstones and their probable influence on migration and accumulation of oil and gas. Am. Assoc. Pet. Geol. Bull. 25:1839-1879.
Walker, R. G. 1984. Shelf and shallow marine sands. P. 141-170. In R. G. Walker (ed.), Facies models. Geol. Assoc. of Canada Reprint Series I. Toronto, Ontario, Canada. 163
Walker, R. G., W. L. Duke, and D. A. Leckie. 1983. Hummocky stratification: Significance of its variable bedding sequences; Discussion. Geol. Soc. Am. Bull. 94:1245-1249.
Walker, T. R. 1960. Carbonate replacement of detrital crystalline silicate minerals as a source of authigenic silica in sedimentary rocks. Geol. Geol. Soc. Am. Bull. 71:145-152.
Walker, T. R. 1962. Reversible nature of chert-carbonate replacement in sedimentary rocks. Geol. Soc. Am. Bull. 73:237-242.
Warner, N. M. 1965. Cementation as a clue to structure, drainage patterns, permeability, and other factors. J. Sed. Pet. 35:797-804.
Washburne, C. W. 1908. Gas fields of the Bighorn Basin, Wyoming. U. S. Geol. Surv. Bull. 340-F.
Washburne, C. W. 1909. Coalfields of the northeast side of the Bighorn Basin, Wyoming and of Bridger, Montana. U. S. Geol. Surv. Bull. 341.
Washington, H. S. 1929. Dahllite from St. Paul's Rocks (Atlantic). Am. Mineralogist Am. Mineralogist 14:369-372.
Waugh, B. 1970. Petrology, provenance and silica diagenesis of the Penrith Sandstone (Lower Permian) of northwest England. J. Sed. Pet. 40:1226-1240.
Weaver, C. E. 1967. Potassium, illite, and the ocean. Geochim. Cosmochim. Acta 31:2181-2196.
Weaver, C. E. 1978. Clay sedimentation facies. P. 159-164. %n R. W. Fairbridge and J. Bourgeois (eds.), The Encyclopedia of Sedimentology. Dowden, Hutchinson and Ross, Stroudsberg, Pa.
Weimer, R. J. 1962. Late Jurassic and Early Cretaceous correlations, southcentral Wyoming and northwestern Colorado. Wyo. Geol. Assoc. Guidebook 17:124-130. 164
Weimer, R. J. 1983. Relation of unconformities, tectonics, and sea level changes, Cretaceous of the Denver Basin and adjacent areas. P 357-376. %n R. J. Weimer (ed.), Mesozoic paleogeography of the west-central United States. Symposium 2. Rocky Mountain Section, S. E. P. M., Denver, Colorado.
Weimer, R. J. 1984. Relation of unconformities, tectonics, and sea-level changes. Cretaceous of Western Interior, U. S. A. P. 7-35. In, AAPG Memoir 36: Interregional unconformities and hydrocarbon accumulation. AAPG, Tulsa, Oklahoma.
Whitlatch, R. B. 1980. Patterns of resource utilization and coexistence in marine intertidal deposit-feeding communities. J. Mar. Res. 38:743-765.
Wilmarth, M. G., Compiler. 1925. Tentative correlation of geologic formations in Wyoming. U. S. Geol. Surv. Comm. Geol. Names, Chart.
Wilson, J. M. 1958. Stratigraphie relations of non-marine Jurassic and pre-Thermopolis Lower Cretaceous strata of northcentral and northeastern Wyoming. Wyo. Geol. Assoc. Guidebook 13:77-79.
Wiltschko, D. V. and J. A. Dorr, Jr. 1983. Timing of deformation in overthrust belt and foreland of Idaho, Wyoming, and Utah. Am. Assoc. Pet. Geol. Bull. 67(8):1304-1322.
Young, R. G. 1960. Dakota Group of Colorado Plateau. Am. Assoc. Pet. Geol. Bull. 44:156-194.
Young, R. G. 1970. Lower Cretaceous of Wyoming and the southern rockies. Wyo. Geol. Assoc. Guidebook 22:147-160.
Ziegler, V. 1917a. The Bryon oil and gas fields. Bighorn County. Wyo. Geol. Surv. Bull. 14:181-207.
Ziegler, V. 1917b. The Oregon Basin oil and gas field. Park County. Wyo. Geol. Surv. Bull. 15:211-242. 165
ACKNOWLEDGEMENTS
This study was made possible through the aid of the
Peace Fellowship Program during the first two years of my graduate study and an Egyptian Mission Fellowship during the rest of the period I stayed in the U. S. A.
I would like to thank and express my deepest gratitude to my major professor, Dr. Carl F. Vondra of Iowa State
University for introducing me to the Bighorn Basin, suggesting the problem, guidance, advice, constructive criticism and support. He has patiently read the manuscript several times and contributed a great deal to its improvement. My sincere thanks and appreciation are also extended to the members of my advisory committee, Drs. John
Lemish, Robert Cody, Thomas E. Fenton and Michael L.
Thompson of Iowa State University, for their interest and advice during this study. Thanks are also due to David Dunn and Nazrul Khandaker of Iowa State University for their assistance in the field and useful comments and suggestions.
I'm deeply indebted to my home country Egypt for endless support. Sincere thanks are due to the Dean and staff of Faculty of Science, Menoufia University, Shebin El-
Kom, Egypt for their support and encouragement. I wish to thank all the Egyptian employees in the Egyptian Mission
Bureau in Washington, D. C. and Cairo, Egypt for their 166 cooperation and support. Sincere thanks to my mother, brothers and sisters for their unending encouragement and moral support. Finally, to my wife, Nadra, and my daughters, Azah, Abir, Aola, and Afaf, whose love, support, encouragement and patience are greatly appreciated, I dedicate this dissertation.