Extensional tectonics of the Cordilleran foreland fold and thrust belt and the Jurassic- Cretaceous Great Valley forearc basin

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EXTENSIONAL TECTONICS OF THE CORDILLERAN FORELAND FOLD AND THRUST BELT AND THE JURASSIC-CRETACEOUS GREAT VALLEY FOREARC BASIN

fcy Kurt Norman Consteniiis

A Dissertation Submitted to the Faculty of the DEPARTMENT OF GEGSdENCES In Partial Ftilfillment of the Requirements For the Degree of DOCTOR OF PHILOSOPHY In the Graduate College THE UNIVERSITY OF ARIZONA

1998 UHI Nximber: 9829328

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THE UNIVERSITY OF ARIZONA ® GRADUATE COLLEGE

As members of the Final Examination Committee $ we certify that we have read the dissertation prepared by Kurt Norman Constenius entitled Extensional Tectonics of the Cordilleran Foreland Fold and

Thrust Belt and the Jurassic-Cretaceous Great Valley Forearc

Basin

and recommend that it be accepted as fulfilling the dissertation requirement for th_e Degree of Doctor of Philosophy

Roy A. Jyohnsori Date f i ^ 1 Peter J. Coney,' -s. Date P Peter G. DeCelles Date

•fA

George H. Dav Date

Final approval and acceptance of this dissertation is contingent upon the candidate's submission of the final copy of the dissertation to the Graduate College.

I hereby certify that I have read this dissertation prepared under my direction and recommend that it be accepted as fulfilling the dissertation requirement. / /i _:£Zj ector Date 3

STATEMENT BY AUTHOR

This dissertation has been submitted in partial fulfillment of requirements for an advanced degree at the University of Arizona and is deposited in the University Library to be made available to borrowers under rules of the library.

Brief quotations from this dissertation are allowable without special permission, provided that accurate acknowledgment of source is made. Requests for permission for extended quotation from or reproduction of this manuscript in whole or in part may be granted by the head of the major department or the Dean of the Graduate College when in his or her judgment the proposed use of material is in the interests of scholarship. In all other instances, however, permission must be obtained from the author.

Signed; 4

ACKNOWLEDGMENTS

No endeavor can succeed without a sound foundation. The basis of my professional and academic achievements is the resolve, encouragement, humor, and inspiration provided by my family. I extend special thanks to Jenny, Matt and Lindsey and to my parents John N. and Leona Constenius. My grandfather John H, Constenius, uncle Raymond Boyce and John Montagne also deserve special thanks for their sound advice. I am grateful to my advisor Roy Johnson for providing guidance and support. Throughout my graduate study at the University of Arizona I have benefited from the teachings, conversations and thoughtful suggestions by Suzanne Baldwin, Susan Beck, Peter Coney, Qem Chase, George Davis, Peter DeCelles, Bill Dickinson, George Gehrels, and Jonathan Patchett. I thank Ned Steme, Jim Coogan and Robert Mueller for many fun and insightful discussions of thrust belt structure. I gratefully acknowledge Tom Vogel, Bill Cambray, LeeAnn Feher and Tim Flood for their interest and significant contributions that they made to my research. I am indebted to the following individuals for assistance they provided: Phil Armstrong, Chris Beard, Rick Bogehold, Mart Bringhurst, Bruce Bryant, Ron Bruhn, William Cobban, Peter Copeland, Brian Currie, Mary Dawson, Carmela Gcirzione, John Groves, Earl Hawkes, Lehi Hintze, Damian Hodkinson, Brian Horton, Pat Jackson, David John, Eric Johnson, Peter Kamp, William Kempner, John Kleist, Tim Lawton, Paul W. Layer, Steve May, Larry Meckel, R. Paul Milne, Gautam Mitra, Susan Olig, Nick Palese, 5

Tim Paulsen, Harold Pierce, Peangta Satarugsa, Alan Satterlee, Alan Tabrum, Kenneth Vogel, Greg Wahlman, John Welsh, Tom Williams, Jim Wilson, Mark Wilson, and Adolph Yonkee. Dan Johnson, Gopal Mohapatra, Carol Revelt, and Craig Steury provided cheerful companionship in the field. I eim especially indebted to Steve Sorenson for his deft programming, computer assistance and many enjoyable discussions. I gratefully acknowledge the guidance and support given by Audrey McCoy, Dave and Ecky Broad, and the late Ted Broad. Seismic data were generously provided by Amoco Production Co., CGG, Exxon Exploration Co., Shell Western Exploration and Production Co., and Texaco Exploration and Production, Inc. Seismic data were reprocessed by Excel Geophysical Services, Inc. Borehole logs and paleontologic data were provided by Amoco Production Co., Canadian Hunter Exploration, Ltd., Chevron, USA, Inc., Exxon Exploration Co., and QC DATA Inc. Interpretation of seismic reflection profiles used Geophysical Micro Computer Application's LOGM software. Landmark's PROMAX software, and Schlvmiberger's Geoquest^^ system. This research was funded by grants from the American Association of Petroleum Geologists Gordon I. Atwater Memorial Fund, Amoco Production Co., Chevron USA, Inc., Colorado Scientific Society, Geological Society of America, Sigma Xi, University of Arizona, and National Science Foimdation grants EAR-9205065 and EAR-9317096. 6

TABLE OF CONTENTS LIST OF ILLUSTRATIONS 9 LIST OF TABLES 10 ABSTRACT 11 CHAPTER 1 INTRODUCTION 13

CHAPTER 2 LATE PALEOGENE EXTENSIONAL COLLAPSE OF THE CORDILLERAN FORELAND FOLD AND THRUST BELT 16

2.1 ABSTRACT 16

2.2 INTRODUCTION 17

2.3 MONTANA DISTURBED BELT AND SOUTHERN CANADIAN ROCKIES 22 2.3.1 Structure of the Lewis Thrust and its Footwall 22 2.3.2 Age of Final Slip on tiie Lewis Thrust 29 2.3.3 Structure of the Flathead Normal Fault 34 2.3.4 Age of Slip on the Flathead Normal Fault 38 2.3.5 Regional Style of Normal Faulting 40

2.4 FOLD AND THRUST BELT OF SOUTHWEST WYOMING- NORTHEAST UTAH 44 2.4.1 Structiure of the Medicine Butte Thrust 47 2.4.2 Age of Final Slip on the Medicine Butte Thrust 48 2.4.3 Structure of the Acocks-Almy Fault System 49 2.4.4 Age of Slip, Acocks-Almy Normal Faults 50 2.4.5 Regional Style of Normal Faulting 55

2.5 TIME-SPACE PATTERNS OF HALF GRABENS, SOUTHERN CANADA-CENTRAL UTAH 58

2.6 LATE CRETACEOUS-PALEOGENE MAGMATISM: A RECORD OF CHANGING PLATE REGIMES 64

2.7 EXTENSIONAL MECHANISMS 72

2.8 STRUCTURAL GENESIS 74

2.9 CONCLUSIONS 76 7

CHAPTER 3 STRUCTURE OF THE DEER CREEK DETACHMENT FAULT SYSTEM AND EVOLUTION OF COTTONWOOD METAMORPHIC CORE COMPLEX, WASATCH MOUNTAINS, UTAH 78

3.1 ABSTRACT 78

3.2 INTRODUCTION 79

3.3 FOLD-THRUST STRUCTURE: STRUCTURAL TEMPLATE FOR COLLAPSE 84

3.4 RECORD OF CHARLESTON-NEBO COLLAPSE 98 3.4.1 Superstructure Collapse and Extensional Triangle Zone 98 3.4.2 New and Previous Concepts 99 3.4.3 Structural Framework of Bowl-Shaped Half-Grabens 104

3.5 DEER CREEK DETACHMENT FAULT 109 3.5.1 Geologic Setting 109 3.5.2 Structural Evidence of Extension 113 3.5.3 Geochronologic Evidence of Footwall Exhumation 141

3.6 COTTONWOOD METAMORPHIC CORE COMPLEX 156 3.6.1 Flexurcil-Isostatic Exhumation & Crustal Welt 156 3.6.2 Bouguer Gravity Model & Discussion 157

3.7 SIGNinCANCE 165

3.8 CONCLUSIONS 167

CHAPTER 4 TECTONIC EVOLUTION OF THE JURASSIC-CRETACEOUS GREAT VALLEY FOREARQ CALIFORNIA; IMPLICATIONS FOR THE FRANCISCAN THRUST-WEDGE HYPOTHESIS 171

4.1 ABSTRACT 171

4.2 INTRODUCTION 172 4.2.1 General 172 4.2.2 Significance of Problem 173 4.2.3 Geologic Setting 179 4.2.4 Paskenta, Elder Creek and Cold Fork Fault Zones 182

4.3 STRATAL RECONSTRUCTIONS OF THE GREAT VALLEY FORE ARC - PASKENTA FAULT ZONE 187 8

4.3.1 Description of Data & Method 187 4.3.2 Interpretation 195

4.4 DOWN-STRUCTURE VIEW OF GREAT VALLEY OUTCROP BELT 207

4.5 COAST RAJSIGE OPfflOLIIE REFLECTIONS: EVIDENCE FOR VOLCANIC RH'TED MARGIN 214

4.6 CONTINUITY OF OPHIOLmC BASEMENT 221 4.6.1 Gravity Modeling - Data and Methods 221 4.6.2 Gravity Modeling Resiilts 223

4.7 CONCLUSIONS 227

CHAPTER 5 REFERENCES 230 5.1.1 References Qted - Chapters 1 & 2 230 5.1.2 References Qted - Chapter 3 250 5.1.3 References Qted - Chapter 4 262 9

LIST OF ILLUSTRATIONS

FIGURE 2.1, Index map of intermontane basins 19 FIGURE 22, Schematic illustration of Cordilleran fold-thrust belt 24 FIGURE 2.3, Index map of Lewis thrust and Kishenehn basin. 26 FIGURE 2.4, Geologic cross section A-A' of Lewis thrust 28 FIGURE 2.5, Geologic cross section B-B' of Kishenehn basin. 37 FIGURE 2.6, Regional cross section R-R' of NW Montana 42 FIGURE 2.7, Geologic index map of SW Wyoming & Utah 46 FIGURE 2 J, Geologic cross section C-C of Medicine Butte thrust 52 FIGURE 2.9, Regionial cross section R2-R2'SW Wyoming & Utah 57 FIGURE 2.10, Qiart showing age distribution of tectonic deposits 63 FIGURE 2.11, Magmatic sweep and initiation of extensional basins 68 FIGURE 2.12, Late Paleogene Cordilleran tectonic elements 70 FIGURE 3.1, Geologic index map north-central Utah 82 FIGURE 3.2, Geologic cross section CN-CN' 87 FIGURE 33, Structural evolution of Charleston-Nebo allochthon 89 FIGURE 3.4, Seismic reflection profile 12 92 FIGURE 3.5, Geologic cross section C-C 102 FIGURE 3.6, Oblique aerial photo Little Diamond Fork half-graben 107 FIGURE 3.7, Geologic index map Cottonwood core complex 112 FIGURE 3.8, Geologic CTOSS section A-A' 116 FIGURE 3.9, Geologic cross section B-B' 118 FIGURE 3.10, Oblique aerial photo of Box Elder Peak anticline 120 FIGURE 3.11, Oblique aerial photo of American Fork Canyon 122 FIGURE 3.12, Hbble Formation paleocurrent directions 125 FIGURE 3.13, Detailed structuri data Deer Creek detachment 130 FIGURE 3.14, Kinematic data Deer Creek detachment 132 FIGURE 3.15, Geologic map Silver Lake detachment 137 FIGURE 3.16, Sketch of Silver Lake drque 139 FIGURE 3.17, U/Pb concordia diagrams 144 FIGURE 3.18, Stunmary of 40Ar/39Ar and AFT data 152 FIGURE 3.19, Bouguer gravity model X-X' 159 FIGURE 320, Cheyenne belt crustal welt 164 FIGURE 4.1, Tectonic models of Great Valley 176 FIGURE 42, Index map study area. Great Valley, and Coast Ranges 181 FIGURE 43, Geologic index map showing seismic profiles and wells 185 FIGURE 4.4, Graph of velocity versus depth for Great Valley Group 190 FIGURE 4.5, Chronostratigraphy of Great Valley monocline 192 FIGURE 4.6, Montage of seismic profiles .194 FIGURE 4.7, Modem forearc examples 197 FIGURE 4.8, Landsat image of Great Valley outcrop belt 206 FIGURE 4.9, Downplunge view of Great Valley outcrop belt 209 FIGURE 4.10, Patterns of seismic reflectivity firom ophiolitic basement 217 FIGURE 4.11, Bouguer gravity model A-A' 225 10

LIST OF TABLES

TABLE 4.1, U/Pb zircon analytical data 145 TABLE 4.2, Summary of WIB geochronologic data 147 11

ABSTRACT

Following cessation of contractional deformation, the Sevier orogenic belt collapsed and spread west during a middle Eocene to middle Miocene (-48-20 Ma) episode of cnistal extension coeval with formation of metamorphic core complexes and regional magmatism. The sedimentary and structural record of this event is a network of half-grabens that extends from southern Canada to at least central Utah. Extensional structures superposed on this fold-thrust belt are rooted in the physical stratigraphy, structural relief and sole faults of preexisting thrust-fold structures. Commonly, the same detachment surfaces were used to accommodate both contractional and extensional deformation. Foreland and hinterland extensional elements of the Cordillera that are normally widely separated are uniquely collocated in central Utah where the thrust belt straddles the Archean-Proterozoic Cheyeime belt crusted suture. Here, the Charleston-Nebo allochthon, an immense leading-edge structural element of the Sevier belt collapsed during late Eocene-middle Miocene time when the sole ttirust was extensionally reactivated by faults of the Deer Creek detachment favdt system and the allochthon was transported at least 5-7 km back to the west. Concurrently, the north margin of the allochthon was warped by flexural-isostatic rise of a Cheyenne belt crustal welt and its footwall was intruded by crustal melts of the Wasatch igneous belt. Collectively, these elements comprise the Cottonwood metamorphic core complex. Extensional processes were also important in the formation of the 12

Jurassic-Cretaceous Great Valley forearc basin. Advocates of a thrust-wedge hypothesis argued that this forearc experienced prolonged Jurassic-Cretaceous contraction and proposed that northwest-southeast-striking fault systems were evidence of a west-dipping blind Great Valley-Frandscan sole thrust and related backthrusts. Based on interpretation of seismic reflection, borehole, map and stratigraphic data, I propose that these faults and associated bedding geometries are folded synsedimentary normal faults and half-grabens. Thus, late-stage diastrophic mechanisms are not required to interpret a forearc that owes much of its present-day bedding architecture to extensional processes coeval with deposition. I 3

CHAPTER 1 INTRODUCTION

The structure and physiography of continental regions that have undergone crustal extension are well documented in highly extended terrains such as the Basia and Range Province in the western United States and the East African Rift Valleys. Less appreciated however, is the role of extension in contractile settings, whether it was concurrent with, or as a successor to, compressional deformation. Studies of ti\e Cordilleran fold and thrust belt have traditionally focused solely on the timing and genesis of contractional structures and dismissed normal faulting and related phenomena as secondary in importance. Yet study of the interrelationships between contractional and extensional structures yields a wealth of information about the physical stratigraphy and structural geometry of a thrust sheet, and most importantly, places limits on the timing of deformation. Likewise, studies of Cordilleran metamorphic core complexes in the hinterland of the orogen often failed to consider the broader reach of extension into the foreland. Evidence presented here supports the hypothesis that the entire Cordilleran orogen collapsed following cessation of Laramide compression, resulting in a chain of half-grabens that parallel the metamorphic core complexes from southern Canada to central Utah and beyond. Similarly, recent literatiu:e regarding the tectonic development of the Great Valley forearc basin and Coast Range subduction complex of California has focused on contractional mechanisms to explain unusual bedding geometries exposed in the Great Valley monocline on the northwest margin of the basin. Unrecognized in these studies of supposed thrust-faults is the 1 4 extensional origin of many of these faxilts and the insights they provide regarding Jurassic-Cretaceous forearc basin evolution. In this dissertation, I address the role of extension in contractional settings by presenting results of: 1) regional study of the Sevier belt collapse, 2) detailed study of extensional structures superposed on the Qiarleston-Nebo allochthon and its association with the Qieyenne belt, and 3) investigating the three-dimensional structural geometry of half-grabens exposed in the Great Valley monocline. Qiapter 2 presents structural evidence and a compilation of space- time patterns of contractional and extensional deformation that bear on the origin of late Paleogene half-grabens in the Sevier belt. Hypotheses are proposed to explain the linkages between extensional structures found in the foreland fold and thrust belt, metamorphic core complexes, regional magmatism, and changes in the mode of North America-Pacific plate convergence. This chapter has been published in the Geological Society of America Bulletin (Constenius, 1996). Chapter 3 focuses on the extensional collapse of the Charleston-Nebo (C-N) allochthon and its interaction with a structurally and magmatically rejuvenated Precambrian crustal suture. This area is situated in an area in which profoimd changes with respect to the cessation of Laramide crustal contraction and initiation of extension occurred. This temporal juxtaposition of major reversal of tectonic styles appear to be a reflection of the Cheyenne belt crustal suture. Based on the structural domains present along the extensionaUy reactivated north margin of the C-N allochthon and the large eimount of extensional exhumation, I propose that this area is a metamorphic 15 core complex. Thus, this area is imique in that normally widely separated tectonic elements of the Cordillera converge at this locality. This chapter was written as a prepublication manuscript co-authored by George Gehrels, Tom Vogel, William Cambray and members of the Wasatch Igneous Belt study team. Chapter 4 investigates the structural geometry of an enigmatic set of northwest-striking faults using seismic reflection, borehole and outcrop data. Study of these faults is important becatise it has been argued that they are evidence of large-magnitude (30-75 km), east-west horizontal shortening, associated with Franciscan thrust-wedge tectonics. Restoration of the depositional geometries of strata in the hanging wall of the faults, down- structure views, stratigraphic data indicating stratal-expansion across the faults and compilation of modem forearc analogs lead to the conclusion that these faults were Jurassic-Cretaceous s5msedimentary normal faults that were uplifted and folded by later deformations. Structural geometries predicted by the thrust-wedge hypothesis are further constrained by characterization of the basement beneath the forearc using seismic, gravity and borehole data. This chapter, co-authored by Roy Johnson, William Dickinson and Tom Williams, has been submitted to the Geological Society of America Bulletin. 16

CHAPTER 2 LATE PALEOGENE EXTENSIONAL COLLAPSE OF THE CORDILLERAN FORELAND FOLD AND THRUST BELT.

2.1 ABSTRACT The Cordilleran foreland fold and thrust belt collapsed and spread to the west during a middle Eocene to early Miocene (ca. 49-20 Ma) episode of crustal extension. The sedimentary and structural record of this event is preserved in a network of half grabens that extend from southern Canada to central Utah. Extensional structures superposed on this allochthonous terrain are rooted to the physical stratigraphy, structural relief and sole faults of preexisting thrust-fold structures. The sole faults dip 3-6° west above an undeformed Precambrian crystalline basement, and accommodated tectonic transport of a thick (up to 20+ km) eastward-tapering hanging wall during regimes of crustal shortening and extension. The chronology of tectonism for the foreland fold and thrust belt is established here by dating latest thrusting and initial normal faulting, and is best defined where thrusts and normal faults are linked by common detachment surfaces. Dated movement on two extensionally reactivated thrusts, the Lewis thrust of northwest Montana and southeast British Coltimbia, and the Medicine Butte thrust of southwest Wyoming and northeast Utah, suggest that the hiatus between the end of crustal shortening in the early or early middle Eocene and the start of extension in the early middle Eocene was brief. Lateral spreading and extensional basin formation in the Cordilleran foreland fold and thrust belt were partly concurrent with formation of metamorphic core complexes and regional magmatism. Conceptually linking 17 extensional processes that were simultaneously deforming both the hinterland and foreland of the late Paleogene Cordilleran orogenic wedge, is accomplished by applying the extensional-wedge Coulomb critical-taper model. The rapid drop in North America-Pacific plate convergence rate and/or steepening of the subducted oceanic slab at ca. 50 Ma resulted in a large reduction in east-west horizontal compressive stress in ttie Cordillera. As a result, the Cordilleran orogenic wedge was left imsupported, and it gravitationally collapsed and horizontally spread west until a new equilibrium was established at ca. 20 Ma. Subsequently, crustal extension and magmatism during the Basin and Range event (ca. 17-0 Ma) overprinted much of this earlier phase of extension.

2^ INTRODUCTION This paper examines a widespread belt of late Paleogene crustal extension that is expressed as a network of half grabens superposed on the Cordilleran foreland fold and thrust belt from southern British Columbia to central Utah (Fig. 2.1). Regional studies of the foreland fold and thrust belt identified the structural style and general ages of extension but considered the half grabens to be of local importance (Armstrong and Oriel, 1965; Bally et al., 1966; Dahlstrom, 1970; Royse et al., 1975; Fermor and Moffat, 1993; Royse, 1993). Detailed studies of individual half grabens in the thrust belt also focused on structure and ages of basin-fill assemblages but failed to consider the development of these structures in the context of an integrated 1 8

Figure 2,1, Index map of intermontane basins of the Cordilleran foreland fold and thrust belt. Distribution of late Paleogene basin-fill (stipple) delineated by surface mapping, exploratory wells and seismic interpretation. Also shown, volcanic fields cited in text (v-shaped stipple). Sources for map: Pardee (1950), Standlee (1982), Love and Christiansen (1985), Fields and others (1985), Whipple, Mudge, and Earhart (1987), Rasmussen (1989), Constenius and others (1989), Bryant (1990), Hanneman and Wideman (1991), Mudge and Earhart (1991), Witkind and Weiss (1991), Bryant (1992), M'Gonigle and Dalrymple (1993), and Janecke and Snee (1993). Basin outlines adapted from Renfro and Feray, (1972). 19

BRiTlSH COLUMBIA ALBERTA llftOH* lU^W iiyw ntw' F CASADA Kishenehn bastn Bearpaw mountains Whitefisft ^

Kaiispell Flathead Lake Htghwood Mountains Great Falls MONTANA

Missoula Eastern Limit Thin-Skinned Thrusting

Heltna Crazy " (jif Buttt Mountains Southwest Montana Province of Basement Involved Thrusts ^ afi'

Boztman (1) Kishenehn basin 12) Tobacco Plains Valley (3) Flathead Valley Horse Prairie (4) South Fork basin (Hunm Horse) i5»Nf— & Medicine Lodge (4b) South Fork basin (While River) basins Mission Valley Swan, Avon-Nevada Creek & Blackfoot Valleys Lincoln Valley Douglas Creek Valley Flint Creek Valley Deer Lodge Valley Helena tc. Townscnd Valleys East-Central Idaho Smith River Valley half grabens Toston Valley Jackson Qarkston basin Madison-Gallatin basin Idaho Falls Norris basin IDAHO Three Forks basin North Boulder basin Jefferson River Valley • Poeatello Lower Ruby River Valley 200 km Beaverhead Valley 1 Divide Trcnch & Melrose Valley Fotvkes & Woodruff French Gulch basin half grabens Big Hole Valley Salmon Valley Grasshopper Valley Horse Pnirie basin Medicine Lodge basin Muddy Creek basin Sage (.reek & Red Rock River basins Great Upper Ruby River Valley 4I<>Nf- Centennial Valley Lemhi tc Birch Creek Valleys East-Central Idaho half grabens Fowkes half fpraben Woodruff half graben Morgan Valley - Huntsville graben Provo • East Canyon half graben Farmington-Wasalch Front basin UTAH Great Salt Lake basin Tibbie half gnben Little Diamond Creek half graben Little Clear Cieek half gnben Santaquin Meadows & Payson Lakes half grabens Juab Valley 390N^ Sanpete Valley Sevier Valley 20 system of extension that affect large parts of the Cordillera (for example, Kuenzi and Fields, 1971; Royse et al., 1975; Sprinkel, 1979; Constenius, 1982). Previous studies generally considered the half grabens to be late Paleogene and Neogene "Basin and Range" phenomenon that developed 10 to 20 million years or more after thrust-fold deformation (for example, Pardee, 1950; McMaimis, 1965; Bally et al., 1966; Constenius, 1982; Lamerson, 1982; Fields et al., 1985). Recently it has been proposed that some of the half grabens encompass a rift zone in Idaho and Montana (Janecke, 1994). The primary goal of this paper is to document the structural setting and ages of contractile and extensiorial deformation that bear on the origin of the late Paleogene half grabei\s. Additionally, data presented here provide evidence for temporal and structural liiJcage of extensioncd structures in the foreland fold and thrust belt with formation of metamorphic core complexes and low-angle detachment faults. The following hypotheses are proposed regarding the transition from contractile to extensional deformation, the ages of extensional deformation, and the genesis of crustal extension in the foreland fold and thrust belt. First, the hiatus between the cessation of crustal shortening and the inception of crustal extension was brief, ~ 1-5 million years. Second, two temporally and tectonically distinct episodes of normal faulting and basin-fill sedimentation are recognized: (1) a late Paleogene episode that was partly coeval with formation of metamorphic core complexes in the hinterland of the Cordillera, and, (2) a Neogene episode related to Basin and Range tectonism. Third, changes in the rate and style of North America-Pacific plate convergence in late Paleogene time resulted in the gravitational collapse of the Cordilleran orogenic wedge. 2 I

Evidence to support these hypotheses comes from tabulation of chronostratigraphic and geochronologic data that describe the timing of latest contractile and the onset and duration of extensional deformation in the foreland fold and thrust belt. The timing of extensional basin formation in the Cordilleran foreland is examined here in both dip and strike directions and with respect to the time-space pattern of regional magmatism. Detailed discussion of two systems of linked thrust and normal fault structures fotind at the northern and southern ends of the half graben network, the Lewis thrust-Kishenehn basin system of northwest Montana and southeast British Columbia, and the Medicine Butte thrust-Fowkes half graben system of southwest Wyoming and northeast Utah (Fig. 2.1), is used to illustrate the timing and structural styles of contractile and extensional deformation. These two basins contain thick sequences of nonmarine, late Paleogene sedimentary rocks that are bounded by listric normal faults that sole into pre­ existing thrust faults, the Lewis and Medicine Butte thrusts, respectively (Bally et al., 1966; McMechan, 1981; Constenius, 1982; Lamerson, 1982). The Lewis and Medicine Butte detachments and related structures were the focus of this investigation because it was possible to construct an excellent record of the ages of thrust and normal displacements on these faults and the structural geometry of these fault systems has been well documented. Three fimdamental relationships establish a tectonic chronology for these areas (Fig. 2.2). First, the end of crustal shortening was determined from dating the youngest thrust(s) by cross-cutting relationships, provenance data, dating of fault gouge, and thermochronologic studies. Second, extensional reactivation of the "youngest" thrust faults was identified where listric normal faults sole 22

dovmward to foinner thrust surfaces. Fineilly, the age of initial crustal extension was determined from dating basin-fill sedimentary imits that were deposited synchronously with displacements on the boimding listric normal fault.

2.3 MONTANA DISTURBED BELT & SOUTHERN CANADIAN ROCKIES 2.3.1 Structure of the Lewis Thrust and its Footwall The Lewis thrust is a classic folded thrust with a strike dimension of 457 km (Mudge and Earhart, 1980) that places a six- to seven-kilometer thick slab of Proterozoic and lower Paleozoic rocks (Ross, 1959; Qiilders, 1964) onto a complexly deformed footwall of Paleozoic and Mesozoic strata (Figs. 2.3 and 2.4). Displacement on parts of the fault have been estimated in the range of 100-115 km (Van der Velden and Cook, 1994; Price, 1988), and may exceed ca. 135 km (Fig. 2.6). A striking feature of this thrust is its prominent, centrally located salient, which protrudes 40-km-northeastward toward the foreland. The Lewis thrust terminates near Mount Kidd, British Columbia in the north and Steamboat Moimtain, Montana in the south, along shear zones in tightly folded Mississippian rocks (Dahlstrom et al., 1962; Mudge and Earhart, 1980). Laterally from these tip-points, the Lewis thrust cut symmetrically down section in the hanging wall, eventually tnmcating Proterozoic rocks. 23

Figure 2.2. Schematic illustration of Cordilleran foreland fold and thrust belt depicting deformational style, principal tectonic elements, and cross cutting relationships used to establish a chronology of tectonism. Timing of thrusting versus normal faulting is definitive, because it is unlikely that the sole fault simultaneously accommodated both shortening and extension. EXTENSIONAL COLLAPSE CORDILLERAN FORELAND FOLD & THRUST BELT

Synextensionat Basw-FIII Struts! Middle Eocene Igneous Rocks (Mlitdle Eocene-Early Miocene) Synorogenic Sedimentary Rocks (Cretaceous-Early Eocene)

Igneous Dikes & Sills

Precambrlan Crystalline Basement Detachment Surface with bath Thrust and Normal Movement

to 25

Figure 2.3. Index map highlighting bilateral symmetry of Lewis thrust and Kishenehn basin, and showing location of units or samples cited for age control. Area formerly marked by greatest degree of late Cretaceous to early Eocene shortening along Lewis thrust fault coincides with area of maximum, middle Eocene to early Miocene extension across Kishenehn basin. Basin-fill of Kishenehn basin. South Fork basin, and Rocky Mountain Trench (Tobacco, Flathead and Swan valleys) outlined by heavy stipple pattern. Also shown, position of Eldorado and Steinbach thrusts, which like Lewis thrust have Belt Supergroup rocks in their hanging walls. Foothills part of fold and thrust belt marked by stipple pattern. Late Maastrichtian and Paleocene foreland-basin deposits delineated by line patterns. Volcanic rocks of the Adel Moimtain field outlined by v-shaped pattern. Mt Kidd Porcupine Hills Fm LEWIS THRUST

"f." i>r; BRITISH COLUMBIA Fernie WkS^^kk.-w

CATE CREEK 'hjtFlATHEAD FAULT

KISHENEHN^ ~ Ji ALBERTA BASIN '^- CANADA U.S.A. MONTANA TOBACCO] VALLEY Cenex 4-liSVr ^ \ L»d«nbnij Willow Creek Fm OBrowning EASTEKN BOUNDARY AULT OF Whitefish ^ FOLD A THRUST BELT FLATHEADMi&::i VALLEY Kalispell Sample of Hoffman et al. (1976)

SOUTH nspvi Flathead FORK Chateau Lake •iXBASlH o

MISSION FAULT Augusta SWAN'. VALLEY SOUTH FORK FAULT SWAN FAULT ^Steamboat ELDORADO & STEINBACH THRUSTS -S9J Ma slU in WiUow Crack Fm "*473 Ma dike culi thrusts ind folds Wolf 27

Figure 2.4. Simplified geologic cross section A-A' of Lewis thrust sheet and Kishenehn basin illustrating structural relief of Flathead duplex, Flathead ramp, and Lewis thrust sheet prior to middle Eocene extension. Preextensional thickness of Lewis thrust sheet was ~ 7-8 km because rocks as young as Jurassic and Cretaceous are in the Lewis hanging wall are preserved beneath Kishenehn Formation. Approximate limits of Flathead ramp designated R -R'. Rock imit abbreviations: Yap-Ymc, formations of Proterozoic Belt Supergroup; C, Cambrian; Pz, Paleozoic vmdifferentiated; Dv, Devonian rocks; PM, Pennsylvanian and Mississippi rocks; JK, Jurassic and Lower Cretaceous rocks; K, Upper Cretaceous rocks. Kilometer! KISHENEHN BASIN LEWiS THRUST SHEET

SlnicturilJUilcl FUlheadr, Ramp tR*R) WILLOW CREEK Ihicknfus of If wis thrust shrrt FM at inception of crustal ntension I ocrnf I - 7'K Am KISHENEHN FM liathead duplet structural relief J km TIIKilSl I Ysh-Ysnl!;i£:l!Si^ flatheail ramp structural relief km p^iAI. DETArHMKNT - WATERTON AUTOCHTHONOUS FLATHEAD CAMBRIAN-CRETACEOUS DUPLEX ROCKS DUPLEX ARCHEAN CRYSTALLINE BASEMENT Kllomeleri

Is) 00 29

The systematic geometry of the lateral ramps reflect changes in the slip and stratigraphic throw on the fault. The symmetric nature of this "scooped- shaped" thrust controlled the extensional geometry of the Kishenehn basin which resulted from inversion along the Lewis thrust (Constenius, 1988). The Lewis thrust has been folded upward ~ 3.0-3.5 km by formation of the Flathead and Waterton duplex zones in the footwall of the thrust (Fig. 2.4) (Bally et al., 1966; Fermor and Moffat, 1993). The geometry of the hanging wall of the Lewis hanging wall is that of a flat in the lower Belt Supergroup strata, whereas the footwall has flat-ramp-flat relations in which the thrust cut upsection from the middle Cambrian to the Upper Cretaceous from southwest-northeast. Pre-erosional thickness of the Lewis thrust sheet is fixed at - 7- 8 km, because Lower Cretaceous strata in the hanging wall were downdropped and preserved beneath basin-fill deposits of the Kishenehn Formation (Fig. 2.4). Additionally, McGimsey (1982), Fermor and Price (1987), and Yin (1993) have documented extensive intraplate thrust-fold structures in lower Belt rocks that locally thicken the Lewis sheet. Collectively, the Flathead ramp and Flathead duplex raised and folded the ~ 7- 8-km-thick Lewis thrust sheet which created an immense structural culmination and locally overthickened the crust.

2.3^ Age of Final Slip on the Lewis Thrust The end of crustal shortening in the southern Canadian Rockies and the Montana disturbed belt is generally considered to be latest Paleocene to early Eocene in age (Bally et al., 1966; Mudge, 1982; Hoffman et al., 1976; Harlan et al., 1988). Cessation of shortening on the Lewis thrust has not been 30

precisely determined because the thrust sheet has been deeply eroded, removing any Late Gretaceous-early Eocene strata that may have overlain or been tnmcated by the fault. Consequently, establishing a date of youngest movement on the thrust requires relative age determination from cross- cutting relationships, indirect geochronologic dating methods, age determination of associated thrusts and footwall-duplex zones formed concurrently with deformation of the Lewis thrust, and provenance studies. The Lewis thrust tnmcates many subsidiary thrusts in the eastern part of the Disturbed Belt (Mudge and Earhart, 1980; Childers, 1965) and, hence, latest movement on the Lewis is relatively young. Interpretation of the Lewis thrust using a sjmchronous thrusting model (Boyer, 1992), suggest that the entire region was undergoing shortening even though only a minor part may have been actively deforming . Furthermore, Boyer (1992) suggested that the youngest imbricate thrusts are part of the Flathead duplex, the site of later superposed crustal extension (i.e., Flathead normal fault on west limb of culmination). Radiometric age determination of thrust fault displacements in the Front Ranges belt of the southern Canadian Rockies give ages of 65-55 Ma for the Lewis thrust, and indicate that age of thrusting for this area spanned from about 100 to 53 Ma (Covey et al., 1994). These thrust-displacement ages were based on K/ Ar dating of illite and illite-smectite clays from fault gouge and are subject to uncertainties as high as 10 m.y. Absolute age determinations from four Montana disturbed belt studies, restrict the terminal thrust-fold event to between ca. 58 Ma and ca. 48 Ma. First, K/ Ar dating (illite-smectite) of low-grade "burial" metamorphism in Cretaceous bentonite beds that 3 1

resulted from stacking of thrust sheets, limits the time of thrusting to between 72 and 56 Ma (Hoffman et al., 1976). Recalculation of some of the youngest dates reported by Hoffman and others (1976) using revised decay constants gives ages of 57.6, 58.3, and 58.4 Ma (Marvin et al., 1980). These are "disturbed" ages rather than "reset" ages, however, which means that the time of burial metamorphism can be no older than the dted apparent ages. Second, in the Wolf Creek area, a sill of quartz monzonite porphyry folded and cut by thrust faults jdelded a K/Ar date (biotite) of 59.6 +/-1.6 Ma (Schmidt, 1978; Whipple et al., 1987). Third, K/Ar dating of intrusive igneous rock bodies that cut or intrude along thrusts in the southern Montana disturbed belt led Mudge (1982) to conclude that most, if not all, deformation in this area occiured during Paleocene time. His conclusion is based in part on dating of an undeformed homblende-monzonite dike that yielded a K/Ar date (hornblende) of 47.5 +/- 1.3 Ma (Schmidt, 1978; Whipple et al., 1987). In the northwest part of the Grazy Moimtains basin, Harlan and others (1988) have published geochronologic and paleomagnetic data regarding the age and timing of imdeformed alkalic igneous rocks that intruded thrust-deformed rocks of the Fort Union Group (Puercan-middle Tiffanian; ca. 66-60 Ma) (Hartman, 1989; Hartman et al., 1989). These data indicate that thrusting occurred prior to the late early or early middle Eocene (ca. 52-48 Ma). The youngest unit truncated by the Lewis thrust is the Campanian Belly River-Two Medicine Formation (ca. 79-74 Ma; Eberth and Ryan, 1992). The youngest sedimentary uiut truncated by thrusts in the footwall of the Lewris thrust (Price, 1986; Mudge and Earhart, 1983) is the Willow Creek Formation, which is broadly dated by vertebrate and invertebrate fossils as 32

Maastrichtian-early Paleocene (ca. 67-63? Ma)(Russel, 1950, 1968; Tozier, 1956). Sills which intrude the Willow Creek Formation are K/Ar dated at 59.6 +/- 1.6 Ma (Schmidt, 1978; Whipple et al., 1987) (Fig. 2.3). Lithologically, this unit consists of about 235-1260 m of nonmarine sandstone, limestone, and shale (Douglas, 1950; Honkala, 1955; Tozier, 1956; Homer, 1989). The lack of coarse clastic rocks in a unit that lies within 20 km of the present erosional trace of the Lewis thrust (Mudge and Earhart, 1983) lead McMannis (1965) to conclude that the main phase of shortening on the Lewis thrust was subsequent to deposition of the Willow Creek Formation. Evidence that shortening continued after deposition of tiie Willow Creek Formation is found in early to late Paleocene (Puercan(?) to early Tiffanian; ca. 65(?)-61 Ma) rocks of ihe Porcupine Hills Formation (Douglas, 1950; Fox, 1990). The Porcupine Hills Formation along the western margin of Alberta syncline has been tilted 2-50O by displacements on underlying thrusts. The yoxmgest rocks at the latitude of the Lewis allochthon that can be linked to thrusts, and the first appearance of cobble-boulder conglomerate in the foreland, are rare, early Eocene sandstone and conglomerate of the Wasatch Formation on the flariks of the Bearpaw Mountains and in the Missouri Breaks diatremes, isolated localities on the Great Plains about 200 km east of the frontal thrusts of the fold-thrust belt (Fig. 2.1) (Heam, 1976; Heam, 1968; Reeves. 1946). Conglomerate clasts of the Wasatch Formation were derived from thrusts exposed to the west such as the Lewis and Eldorado thrusts, and consist of 50-80% argillite and quartzite of Belt Supergroup, 1-5% Paleozoic rocks, 20-40% Late Cretaceous-Paleocene fine-grained and porphyritic volcanic rocks (Elkhom and Adel Mountain volcanic rocks?), and 33 less than 1% chert and conglomerate (Heam et al., 1964). The average grain- size of the conglomerate clasts and the percentage of limestone clasts decreases from west to east across the Bearpaw moimtains - an indication of eastwcird transport from a western source area (Heam et al., 1964; Bryant et al., 1960). The Wasatch Formation has been assigned an early Eocene age (Wasatchian; ca. 57-54 Ma) based on flora and vertebrate fossils (Marvin et al., 1980; Brown and Pecora, 1949) and is overlain with angular unconformity by extrusive rocks of the Bearpaw Mountains volcanic field (Heam, 1976) which have been isotopically and paleobotaniccdly dated as late early-early middle Eocene (ca. 54-50 Ma) (Marvin et al., 1980; Wing and Greenwood, 1993). The Missouri Breaks diatremes are early middle Eocene in age (ca. 52-47 Ma) (Marvin et al., 1980). Similar stratigraphic relationships are found in the Highwood Moimtains, 100 km southwest of the Bearpaw Moimtains, where early- middle Eocene volcanic rocks (ca. 53-50 Ma) overlie and intrude coarse alluvial-fan conglomerate of the Wasatch Formation (Marvin et al., 1980; O'Brien, 1991). Wasatch conglomerate in the Highwood Mountains consists of clasts of Archean granitic gneiss, Cambrian quartzite, and Paleozoic limestone and dolomite, indicating a different source area than related conglomerate in the Bearpaw Mountains. However, both units represent syntectonic sedimentation associated with the unroofing of Laramide foreland and thrust-fold stmctures, respectively. The Wasatch Formation is the last unit related to thrust-fold shortening in the northem Great Plains, and the early to middle Eocene igneous rocks that overlie and intrude this unit signal the onset of Mid-Cenozoic crustal extension and magmatism in 34

the northern Cordillera.

2.3.3 Structure of the Flathead Normal Fault The Kishenehn basin is situated southwest of the Lewis thrust salient (Fig. 2.3) as a narrow, asymmetric graben 150 km long and 2 to 13 km wide. The Flathead listric normal fault system, the master structure that controlled basin origin, borders the Kishenehn basin to the northeast. Estimates of maximum slip on the Flathead fault are on the order of 15 km (Constenius, 1988). The southwest basin margin is either boimded by faults antithetic to the Flathead system, or it is onlapped by Kishenehn basin strata. Subsidence along these faults created an asymmetric graben containing up to 3,400 m of nonmarine, late Paleogene sedimentary rocks of the Kishenehn Formation. The structural position of the Kishenehn basin with respect to the Lewis thrust salient reflects the common fault stirface that accommodated both shortening and extension (Constenius, 1988; McMechan, 1981). The termini of the basin and the longitudinal extent of the Flathead fault system roughly coincide with the re-entrants of the Lewis thrust salient (Fig. 2.3). In addition, slip on the Flathead and Lewis faults is bilaterally symmetrical, and their centerlines of symmetry and inferred lod of maximimi displacement coincide (Constenius, 1982; McMechan, 1981). Ultimately, theses relationships are related to the symmetry and dimensions of the Lewis footwall ramp (Flathead ramp; Boberg, 1993) which cuts upsection across ~ 2.5 km of middle Cambrian to Lower Cretaceous strata (Figs. 2.4, 2.5, and 2.6) (Bally et al., 1966; Fermor and Moffat, 1993). Critical relationships identified in figures 2.4 and 2.6 are: 35

(1) The Flathead fault is superposed on the west flank of a major structural culmination comprising two overlapping tectonic elements, the Flathead duplex and the Flattiead ramp (Fritts and Klipping, 1987), (2) the Flathead faiilt is a listric normal fault that merges with the reactivated segment of the Lewis thrust. The fault dips ~ 40-50O southwest near the surface and flattens at depth, and is layer-parallel in the basal C

Figure 2.5. Geologic cross section B-B' of southern Kishenehn basin showing listric normal faults of Flathead fault system (Roosevelt and Blacktail faults) soling into reactivated part of Lewis thrust and associated rotation of hanging- wall strata. Kishenehn strata display a gradual flattening of dip upsection and thicken toward Roosevelt fault, a manifestation of concurrent sedimentation and slip on a curved fault surface. Stratigraphic position of middle Eocene (Uintan) fossil localities and dated tephra deposit are indicated. Rock unit abbreviations: Yh-Ymc, formations of Proterozoic Belt Supergroup; Ted, Teem, Tccu and Tp, Coal Creek (lower, middle and upper sequences) and Pinchot members of Kishenehn Formation. to

Cr/sial Creek

Middle Flathead River

tft a- 5 a*a a to ac«

_ Wolftail A' Mountain § ^ 38

2.3.4 Age of Slip on the Flathead Normal Fault The synextensional nature of the Kishenehn Formation has been established using study of sedimentary structures, fades relationships, provenance, paleocurrent directions, and stratal growth geometries (Constenius, 1981,1982,1989; McMechan and Price, 1980; McMechan, 1981). Sjmextensional sedimentary sequences display stratal growth relationships, a systematic thickening of strata toward the basin-bounding listric normal fault and a gradual flattening of dip in successively younger units (Dahlstrom, 1970; McMechan and Price, 1980). The strong rotational control on sedimentation associated with listric normal faulting results in a half graben or asjonmetrical graben with a wedge-shaped sedimentary prism. Strata of the Coal Creek member of the Kishenehn Formation dip 50° NE at the base and progressively decrease to 32° NE near the top of the unit (Fig, 2.5). This indicates that displacement and rotation along the Roosevelt and BlacktaU fault segments of the Flathead fault system was synchronous with sedimentation. Many investigators have used the age of the Kishenehn Formation as the time of initial crustal extension in the region and, consequently, as an upper time limit of contractile deformation (e.g., McMaimis, 1965; Bally et al., 1966; McMechan, 1981; Constenius, 1982). Previous age determinations of the BCishenehn Formation, which relied on dating fossil mammals, mollusks, leaves and pollen from exposures in Canada (Russel, 1954, 1964; Hopkins and Sweet, 1976, McMechan, 1981), concluded that the unit was late Eocene to early Oligocene in age. Recent discoveries of mammal fossils from the Coal Creek member of the Kishenehn Formation suggest a middle Eocene age 39

(Uintan; ca. 48-42 Ma) (M. R. Dawson and A. R. Tabrum, pers. commun., 1993). Isotopic analysis of a tephra from the Coal Creek member also established a middle Eocene age for this unit. Single-crystal laser fusion 40Ar/39Ar dating of 12 biotite grains from this tephra resulted in an age of 46.2 +/- 0.4 Ma (R. C. Walter, written commun., 1990). Fission-track analysis of seven zircon grains from this tephra yield a date of 43.5 +/- 4.9 Ma (C. W. Naeser, written commun., 1990; revised from 33.2 +/-1.5 Ma reported by Constenius et al., 1989). Approximately 1150 m and 1450 m of exposed section imderlie the dated tephra and mammal beds, respectively. Using assimied sedimentation rates on the order of 500 m/m.y. (McMechan, 1981), age estimates of the lower exposures of the Coal Creek would be 2.0 to 2.5 Ma older than the dated units, or roughly 48-49 Ma; early Uintan to Bridgerian. Palynological analysis of a composite cuttings sample from the Cenex #4-13 Ladenburg well (Fig. 3) foimd a small number of late Paleocene-middle Eocene specimens in Kishenehn cavings (Bujak Davies Group, written commim., 1989). Synextensional sedimentation continued imtil the late Oligocene-early Miocene (Arikareean; ca. 25-21 Ma) (H. G. Pierce, written commun., 1993; M. R. Dawson, pers. commun., 1994). The Kishenehn Formation is overledn along a pronounced angular imconformity by a thin veneer (0-100 m) of late Neogene alluvial gravels and glacial detritus, and no middle Miocene to Pliocene deposits have been found. There is little evidence of Quaternary- Recent movement of the basin-boxmding faults with the exception of the mountain-front fault and cissodated triangular facets at Nyack Flats. Hence, subsequent to the middle Eocene-Oligocene episode of extension and 40 sedimentation, the basin has been comparatively qiiiescent and was Icirgely unaffected by Bcisin and I^ge extension (ca. 17-0 Ma).

2.3.5 Regional Style of Normal Faulting The Flathead fault defines the eastern boundary of a - 180-km-wide belt of extension in the Cordilleran foreland fold and thrust belt (Fig. 2.6). Listric normal faults boimding the Kishenehn basin. South Fork basin, southern Rocky Mountain Trench (Flathead Valley), and niunerous other normal faults sole into the extensionally reactivated regional detachment fault. Hanging wall rocks above this regional detachment constitute a westward thickening extensional wedge with a tip delineated by the Flathead favdt and a maximimv thickness of about 20 km foimd west of the Purcell anticlinorium. Total extension across tiiis belt has been estimated at ~ 25 km (McMechan and Price, 1984; Harrison, 1988), but may exceed ~ 30 km because displacements on listric normal faults bounding the Kishenehn basin. South Fork basin and southern Rocky Moimtain Trench are - 15 km, ~ 7 km, and ~ 10 km, respectively (Constenius, 1981, 1988; Van der Velden and Cook, 1994). Numerous small normal faults that have displacements in the range of 0-100 m, that are not resolvable seismically or depicted on regional maps or cross sections collectively contribute significantly to toteil extension. 41

Figure 2.6. Simplified regional geologic cross section Ri-Ri'of northwest Montana based on results of seismic reflection and refraction profiles, Bouguer gravity data, well control and balanced cross-sections (Bally et al., 1966; Harris, 1985; Fritts and Klipping, 1987; Constenius, 1981,1988; Boberg et al., 1989; Yoos et al., 1991; Harrison et al., 1992; Sears and Buckley, 1993; Van der Velden and Cook, 1994). Hanging wall rocks above west-dipping regional detachment constitute a westward ttiickening extensional wedge with a tip delineated by Flathead fault and -20 km maximum thickness west of Purcell anticlinoritmi. The Archean crystalline basement (fine-line stipple) has remained undeformed in spite of considerable supracrustal shortening and extension. Rocks of lower Proterozoic Belt Supergroup shaded (Prichard Formation and equivalents), middle and upper Belt rocks (Ravalli-Missoula Groups) and Phanerozoic rocks unshaded. Cretaceous intrusive rocks line stipple, late Paleogene basin-fill lightiy shaded. Sense of motion on thrusts and extensionally reactivated thrusts indicated (arrows), other normal faults unmarked. R1 Purcell Jocky Kishettehtt Dry Ck anticlinorium Rl' Lightning Ck Mountain Lewis Km Hope slock Moule / Ravalli-Missoula Area-Marathon Trench South Fork thrust fault / thrust^/ Cpa " Gibhs basin Flathtad sheet Pinkham thrust Hefty thrust /"""

FPRICWFI ;' ; >-( i; Archt^^n ^pf^aUine batitmtnt'

• • :

N>4^ 43

Interpretations of seismic reflection profiles, Bouguer gravity data, well control and balanced cross sections in northwest Montana and southern British Columbia indicate that the regional detachment fault dips west at -3® and that the imderlying Archean crystalline basement is undeformed (Bally et al., 1966; Harris, 1985; Fritts and Klipping, 1987; Boberg et al., 1989; Yoos et al., 1991; Sears and Buckley, 1993; Van der Velden and Cook, 1994). The regional detachment and crystalline basement may be warped upward in response to displacements on the Hope fault (Fillipone and Yin, 1994), or imbricate slices of crystalline basement are stacked above the gently west-dipping regional detachment (Harrison et al., 1992) (Fig. 2.6). The dip of the regional detachment may locally steepen to 20-30O under the west limb of the Purcell anticlinoritim that is interpreted here as a fault-bend fold above autochthonous Belt Supergroup rocks. The ramp segment of the regional detachment fault dips west at -17° above autochthonous Belt Supergroup rocks in southern Canada (Van der Velden, and Cook, 1994). One himdred kilometers west of the Flathead fault, middle and lower crustal elements constituting the Priest River complex are juxtaposed with supracrustal rocks across the Newport and Purcell Trench faults. Crustal attenuation and tectonic denudation associated with these faults took place from about 55-42 Ma, and estimates of total horizontal extension range from 35 to 68 km (Harms and Price, 1992; R.A. Price, written commun., 1995). Syntectonic volcanic and sedimentary rocks downdropped and preserved in the hanging wall of the Newport fault yield isotopic ages of ca. 51 Ma and microfloral ages of early to middle Eocene, which are partly coeval with 44

Kishenehn sedimentation. Metamorphic core complexes in the Omineca Belt, of which the Priest River complex is a part, sustained extension from about 60 to 47 Ma (Parrish et al., 1988). In the southern Omineca belt, cross cutting relationships of intrusive rocks, thrusts, and normal faults suggest that extension overlapped with or closely followed shortening. This change in tectonic regime took place in late Paleocene to early Eocene time (58 +/- 1 Ma) (Carr, 1992).

2.4 FOLD AND THRUST BELT OF SOUTHWEST WYOMING - NORTHEAST UTAH

In southwest Wyoming and northeast Utah a similar suite of linked contractile and extensional structures exists; the Medicine Butte thrust, the Acocks-Ahny listric normal fault system and the Fowkes half graben (Figs. 2.1 and 2.7). Their dimensions, however, are an order of magnitude smaller than the Lewis thrust-Kishenehn basin. Analysis of the deformational history is more straight-forward in this area because syntectonic sedimentary rocks that record both phases of deformation are preserved in juxtaposition with thrust-fold and extensional structures. These relationships have been defined by detailed surface mapping, biostratigraphic analysis, and extertsive borehole and seismic data. In addition to the Fowkes half graben, there are several other late Paleogene half grabens in this region, the bounding faults of which 45

Figure 2.7, Geologic index map of southwest Wyoming and northeast and central Utah highlighting late Paleogene grabens and major structural elements of the Cordilleran foreland fold and thrust belt. Structures and basin outlines adapted from Hintze (1980), Love and Christiansen (1985), Witkind and Weiss (1991), and Bryant (1992). 46

1130W Idaho 1120W mow Qc«r Lake

Woodruff half grabea WDF-1

Fawkes

Marfan VafUy HuntsvUic R2I

Amoco State Uiah'L Bridge 1130W Porcupine Wyomine iitpw 410N-i- Ridge —i " i 41 East Caityan half trabett

Great Salt Lake basins SaltQty Uk

Farmington' Wasatefi Front basia

Traoerse Rante volcanic fiel Keetley & Park C/ly Vbble ooleanie fields Cettaawcod arch €r half grabea Sale iVnsatch intrustoe belt 50 km Ptovo Utth LaJce ^I^Fowkw Foniutioo fc Norwood Tuff Meadows Utile Paysoa Lakes Diamond mow half grabens Creek half grnben Norwood Toff-Moroni FormaUon O b cQUtvslcnlt Ute eocene01lgo«ne Inieoof rock« Charleston'Nebo (Volanlc it *0ilcanldMtie • ttippic. thrust IntnitWe - bljek) little Clear Crttk Tintic Moaatatas half graben rTYcambtian rocks volcanic field X f Sanpete Valletf Thrust fault Nornul rjuit fuab Valley 47

are interpreted to sole into pre-existing thrusts. However, west of the Wasatch fault, these half grabens have been overprinted by later Basin and Range extension. Two distinct phases of extension are exemplified by interpretation of seismic and borehole data from the Great Salt Lake that reveal a late Paleogene half graben buried imder several kilometers of Neogene basin-fill.

2.4.1 Structure of the Medicine Butte Thrust The Medidne Butte thrust is a fronted footwall imbricate of the Crawford thrust that is situated above the major footwall ramp and associated fault-bend fold of ttie Absaroka thrust (Coogan, 1992; DeCeUes, 1994). In the footwall ramp ttie Absaroka thrust cuts from Cambrian through to Lower Cretaceous strata. The resultant fault-bend fold in the hanging wall has about 5.0 to 5.5 km of structural relief and is cored by Lower Paleozoic rocks (Lamerson, 1982; Piatt and Royse, 1989). Slip on the Medicine Butte thrust is difficult to determine, but is at least three to four kilometers. Siuface and subsiuface data indicate that at its leading-edge the thrust consists of a zone of imbricate thrust slices of Jurassic Preuss Formation and Upper Jurassic - Lower Cretaceous Gaimet Group. The oldest rock vmit in the Medicine Butte thrust sheet is the evaporite unit of the Preuss Formation. The Preuss evaporite unit is one of the main detachment surfaces, not only for the Medicine Butte thrust, but also the Bridger Hill thrust, the Absaroka thrust and its many imbricates, and the Acocks and Almy listric normal faults (Royse, 1983; Lamerson, 1982). 48

2.4^ Age of Final Slip on the Medicine Butte Thrust The Medidne Butte thrust fault truncated and displaced units as young as the late Paleocene part of the Evanston Formation. Exposxires of even younger units such as the basal conglomerate, lower member and Main Body of the early Eocene Wasatch Formation in the Medidne Butte footwall have sustained thrust-related folding and internal deformation. Exposures of near- vertical or overturned Evanston Formation and/or basal conglomerate member of the Wasatch Formation can be found along the trace of the thrust (Veatch, 1907; Lamerson, 1982; Bryant, 1990). Footwall rocks consisting of the Main Body of the Wasatch and Green River formations dip from 62° to 17° away from the toe of the Medicine Butte thrust. Data relevant to the age of the Evanston and Wasatch formations is included in studies by Gazin (1952; 1956; 1962; 1969), Oriel and Tracey (1970), Jacobson and Nichols (1982), Lamerson (1982), and Nichols and Bryant (1990); a synthesis of their work follows. The upper Evanston has been assigned a late Paleogene (Torrejonian- Tiffanian; ca. 63-59 Ma) age based on vertebrate fauna, leaves and palynomorphs. The age of the basal conglomerate member of the Wasatch is indeterminate. Oriel and Tracey (1970) assigned the imit an early Eocene age based on a single gastropod fossil, whereas Nichols and Bryant (1990) consider its age to be late Paleocene. Rocks of the lower member imconformably overlie the basal conglomerate member, and the fossil fauna and flora suggest an early Eocene age. The main body of the Wasatch contains an extensive vertebrate fauna with early to middle early Eocene a^iiuties (i.e.. Gray Bull and Lysite ages of the Wasatchian; ca. 57-53 Ma). The remaining three 49

members of the Wasatch, the sandstone tongue, the mudstone tongue and the Ttmp member are not preserved in direct contact with the Medicine Butte thrust but do provide a record of early to early middle(?) Eocene contractile deformation. Hurst and Steidtmann (1986) concluded that the three discrete belts of coarse clastic rocks of the Tunp member are synorogenic deposits generated by the uplift of the Absaroka thrust sheet (i.e., passive uplift and rotation over the Hogsback footwall ramp), and last displacements on the Tunp and Crawford tlirusts. Although fossils have not been found in the Tunp member, stratigraphic relationships indicate that its age is equivalent to the whole of the Wasatch Formation and, therefore, it is mainly early Eocene (ca, 57-51 Ma) in age. The upper part of the Tunp may be latest early to middle Eocene based on stratigraphic correlation and dating of gastropods in stratigraphically equivalent beds in the Green River Formation. Thus, the timing of the Medicine Butte thrust is very young and coeval with movement on the Hogsback and Ttmp thrusts. Consequently, this area provides evidence that the region was subjected to regional shortening through early Eocene and possibly early middle Eocene (ca. 57-51 Ma) time.

2.4.3 Structure of the Acocks-Almy Fault System The Acocks-Almy listric normal fault system has reactivated the segment of the Medicine Butte thrust that is superposed on the west limb of the Absaroka fault-bend fold. Lamerson (1982, p. 336) noted that "nearly all the extension in the southern Fossil Basin by normal faults such as the Acocks-Almy system, is concentrated on the hanging wall of the Medicine Butte thrust." About three kilometers of retrograde displacement has teiken 50 place on the Medicine Butte thrust plane. Downdropping and rotation of the Medicine Butte hanging wall along the Acocks-Almy listric fault system created a network of half grabens in which up to 1500 m of middle Eocene Fowkes Formation and at least ~600 m of the Norwood Tuff were deposited and preserved (Figs. 2.8 and 2.9). Interpretation of borehole data, field evidence and seismic data support the contention that Fowkes sedimentation was coevcil with Acocks-Almy normal faulting (Lamerson, 1982, p. 332-334). Principal indications of S5niextensional sedimentation in the Fowkes half graben are: (1) the Fowkes Formation is foimd only in the hanging wall of these faults, (2) thickening of units into the fault and in certain areas flattening of dip in successively younger beds (Figs. 2.8 and 2.9), and (3) development of unconformities within the Fowkes Formation limited to the western margin of the half graben. Parts of the Fowkes half graben where the sjmextensional units lack stratal growth geometries but are tilted may indicate considerable displacement and rotation after deposition of the Fowkes Formation, dissolution and collapse of the Preuss evaporite unit in the footwall of the Acocks-Almy fault system, displacements of diapir-like bodies of Preuss evaporite, and variations in the dip of the fault surface (Matos, 1993; EUis and McQay, 1988), or a combination tiiereof. The Preuss evaporite unit in tile footwall of the Acocks-Almy fault system attains thicknesses as much as 1.2 km (Lamerson, 1982).

2.4.4 Age of slip Acocks-Almy normal faults The Fowkes Formation contains the first sedimentary record of Figure 2.8. Geologic cross section C-C showing Medicine Butte thrust, Acocks normal fault, and Fowkes half graben above complexly deformed Absaroka thrust hanging wall (modified from R.E. Mueller, Amoco Production Company, 1987). Note progressive unconformities in footwall (Evanston and Wasatch formations) of Medicine Butte thrust and soling of faults into evaporite of Jurassic Preuss Formation. Rock imit abbreviations: Ob-Jtc, formations ranging from Ordovidan Big Horn Dolomite to Jurassic Twin Creek Limestone; Jpev, evaporite unit of Jurassic Preuss Formation; Jsp-Kf, formations ranging from Jurassic Stump-Preuss Formations to Cretaceous Frontier Formation; Kev, Maastrichtian Evanston Formation; Tev, Paleocene Evanston Formation; Tw, late Paleocene and early Eocene Wasatch Formation; Tf, middle Eocene Fowkes Formation; and Tnt, late Eocene Norwood Tuff. Fowkes half graben - Medicine Butte thrust Anschutz Ranch Field Anschutz Ranch East ANSCHUTZ Field Northwest JJ-I ^.K.£. Southeast ANSCHUTZ ANSCHUTZ A.R.E. VVII-OI , THOUSAND PEAKS Meters 3-1 E't 28 ta2if-2 Meters 3000 4 JOOO Acocks Nonnal Fault

Tw-Tev-KTei Medicine Butte^ Thrust 1000

• Km

- 'tm

Upper Cretaceous Rocks . -3000

Upper Cretaceous Rocks « '400Q

•5000

to 53

extension in this area. The Fowkes vinconformably overlies the Wasatch Formation and consists of up to 1500 m of dominantly tuffaceoiis mudstone, sandstone and conglomerate and siliceotis limestone (Nelson, 1973). A middle to late Eocene age for the Fowkes Formation was assigned by Oriel and Tracey (1970) based on fossil gastropods, leaves, and a hornblende K/Ar date of 48.2 +/-1.5 Ma (recalculated). More recent vertebrate paleontologic and radiometric work by Nelson (1973; 1974; 1979), has established that the lower Fowkes Formeition is early middle Eocene in age (Bridgerian; ca. 49-48 Ma). The biotite K/Ar age from a tephra in the Fowkes Formation in Nelson's study was 49.1 +/-1.9 Ma (recalculated using critical table; Dalrymple, 1992). In the southern part of the Fowkes half graben, on Porcupine Ridge, the Fowkes Formation is overlain by late Eocene to late Oligocene synextensional deposits of the Norwood Tuff (Bryant, 1990) (Fig. 2.8). Rocks of the Norwood Tuff have not been dated at Porcupine Ridge, but the Norwood Tuff and its stratigraphic equivalent, the Moroni Formation, are found in association with several other half grabens in northeast and central Utah (for example. East Canyon, Morgan VaJley-Himtsville, Great Salt Lake, Tibbie, Little Diamond Creek; Figs. 2.1 and 2.7). The Norwood Tuff and Moroni Formation, which ranges in thickness from about 1.0 to 4.5 km, consists of tuff, volcaniclastic sandstone and conglomerate, lahars, a few thin flow breccias, interbedded with conglomerate containing sedimentary clasts (Bryant et al., 1989; Witkind and Marvin, 1989; Bryant, 1990). Sjmextensional stratal growth geometries of Norwood Tuff-Moroni Formation deposits in these half grabens indicate that sedimentation was concurrent with normal faulting in late Eocene to late Oligocene time (Royse, 1983; Hopkins and 54

Bruhn, 1983; Riess, 1985; Houghton, 1986; Bryant et al,, 1989; Constenius, 1995). Collectively, isotopic age determinations from these basin-fill deposits suggest that the Norwood Tuff-Moroni Formation is late Eocene to late Oligocene (ca. 39-27 Ma) in age (Crittenden and Sorensen, 1985; Van Horn and Crittenden, 1987; Bryant et al., 1989; Witkind and Marvin, 1989). Mammal fossils from the Norwood Tuff are rare but indicate a late Eocene (Duchesnean and Chadronian; ca. 41-34 Ma age range) age for these rocks (Nelson, 1971). Recent discovery of the camel fossil, Blickomylus near B. galushai, from the upper part of the Moroni Formation may extend the time- range of this formation to the late early Miocene (Hemingfordian; ca. 21-16 Ma) (M. R. Dawson, pers. commun., 1994). The Keetley volcanic field, which was the source of much of the volcaniclastic detritus in the Norwood Tuff, has been paleontologically dated as late Eocene (Chadronian; ca. 37-34 Ma) near the base and isotopic dates range from about 39-33 Ma (Best et al., 1968; Nelson, 1977; Bryant et al., 1989; Bryant, 1990). Neighboring volcanic fields, such as the Tintic Moxmtains, Park Qty and Traverse Range volcanic fields, which are also thought to be sources of sediment for the Norwood Tuff and Moroni Formation, range in age from ca. 40-28 Ma (Laughlin et al., 1969; Nelson, 1971; Morris and Lovering, 1979; Crittenden et al., 1973; Bromfield et al., 1977; Bryant et al., 1989). Physiographically, some exposures of basin-fill in the Fowkes half graben, such as those found on Porcupine Ridge, are inverted with respect to the modem drainage basiii by as much as 200-300 m implying that this basin is no longer an actively subsiding. In low-lying areas cdong the Bear River drainage, the Fowkes Formation is overlain by ordy a thin mantle of 55

Quaternary sediments. Hence, the Fowkes half graben was structurally active from early middle Eocene to late Oligocene time (ca. 49-27 Ma). However, fault scarps associated with Quaternary to Recent extensional reactivation of the Hogsback and Absaroka thrusts have been mapped 15 to 25 km east of the Fowkes half graben (West, 1993).

2.4.5 Regional style of normal faulting Extensional reactivation of the Medicine Butte thrust and formation of the Fowkes half graben were, in part, synchronous with the late Eocene to late Oligocene development of the East Canyon, Woodruff, Morgan Valley- Huntsville, and Great Sedt Lake half graben. These structures are superposed on the Wasatch culmination, a thrust structure that had -10 km of structural relief (Coogan, 1992; Yonkee, 1992; Royse, 1993; DeCelles, 1994)(Figs. 2.7 and 2.9). The faults which boimd the Morgan VaUey-Huntsville, Farmington- Wasatdi Front and Great Salt Lake half graben and the Salt Lake salient are interpreted as west-dipping listric normal faults that sole into the basal Cambrian footwall detachment that had been the master fault for earlier thrusting (Royse et al., 1975; Coogan, 1992). The East Canyon normal fault is an east-dipping fault related to extensional rejuvenation of the East Canyon- Crawford-Medidne Butte thrust complex (DeCelles, 1994). The basal Cambrian detachment dips -3-6° west and the imderlying crystalline basement is tmdeformed (Royse et al., 1975; Lamerson, 1982). Above 56

Figure 2.9. Simplified regional geologic cross section R2-R2' of southwest Wyoming and northeast Utah that emphasizes late Paleogene extensioncd structures and growth-fault depositional patterns (modified from Coogan, 1992; Lamerson, 1982; Wilson et al., 1986). Fowkes half-graben (FWK-1, nonmigrated seismic profile): lower strata of Fowkes Formation exhibit gradual flattening of dip upsection and thicken toward Acocks fault and imconformably overlie strata of Wasatch Formation. Woodruff half-graben (WDF-1, nonmigrated seismic profile): strata of Fowkes Formation display flattening of dip upsection and thicken toward basin-bounding fault. Outcrops of Fowkes Formation on west-side of half graben dip ~10-15O east (Dover, 1985). Wasatch Formation shows stratal onlap on Paleozoic rocks of Crawford thrust. Great Salt Lake basin (GSL-3, nonmigrated seismic profile): seismic profile images flat to gentie apparent dip of Neogene strata, regional Hemingfordian-age(?) angular unconformity separatirig late Paleogene from Neogene strata, late Paleogene half graben buried beneath ~2750 m of

Neogene basin-fill, and north-dipping (~35^) listric normal fault bounding half graben to south. Late Paleogene and Neogene basin-fill deposits delineated by marked dip divergence and lithologic changes in the Amoco State Utah-L Bridge well (projected ~1700 from northecist) (Bortz et al., 1985; Bryant et al., 1989). Analysis of dipmeter data indicate Neogene strata dip 5°- 120 east to northeast, whereeis, late Paleogene beds dip 15-35° south to southwest. ^Ar/39Ar plagioclase ages bracket xmconformity and establish ages for rocks in buried half graben: 10-11 Ma (10.3 +/- 1.0 Ma zircon fission- track) and 38.5 +/- 0.2 plateau (29.9 +/-1.3 Ma zircon fission-track) at depths of 2,721 and 3,674 m, respectively (Bryant et al., 1989). GSL-3 WDF-l Great Salt Lake basin VVrsI Woodruff halfgrahen iMit

Northwfut Southeast ®®|~ Amoco Utah-L A .^uat^rnary ^luvium 00

Q ./ Fowkes Wasatch Vi Neogene rocks Fm .10 -40j^r/^9Ar 10-11 IMa EWK-1 ou 0.5 km 10.3 +/- 1.0 Ma Northwrat Fowkes half grnben ^•0 km Listric normal 00 ' WasatchFm 40/{r/39/{r 38,5 +/. 0.2 Ma 10.5 fcm » o«: 01 Fowkes SJI km Norwood Tuff Vj Evanston ^\ eqv., Fm 0.5 km Listric normal fault •Acocks normal 2.0 km Woodruff fault R2 Wasatch Culmination half graben Morgan Valley .. .. . Fowkes Great Salt Lake basin Wasatch fault Huntsville graben -^""buiH l-'f g">"" -A S SI

a. 10 u V. Reactivated sal^detachment JHMH ^ Q ... crystalline basement 58

this regional detachment, hanging wall rocks form an extensional wedge with a maximum thickness of -10-11 km at the Wasatch culmination. Therefore, beginning in middle Eocene and continuing through late Oligocene time, the -10-11 km thick hanging wall above the basal Cambrian detachment in this region was displaced horizontally to the west. Total net horizontal extension in the hanging wall of the sole fault is on the order of 8- 10 km for the area east of the Wasatch fault (Royse, 1993) and may be as much as 20 km for the region shown in Figure 2.9. Similarly, the half grabens superposed on the Qiarleston-Nebo allochthon record a phase of late Eocene to early Miocene tectonism, in which the sole thrust was extensionally reactivated and the -5-6 km thick allochthon was displaced about 5-7 km to the west (Royse, 1983; Reiss, 1985; Houghton, 1986; Constenius, 1995).

2.5 TIME-SPACE PATTERNS OF HALF-GRABENS, SOUTHERN CANADA-CENTRAL UTAH

The Kishenehn basin and Fowkes half graben are elements of a network of late Paleogene half grabens that extends from southern British Columbia to central Utah and is superposed on the Cordilleran foreland fold and thrust belt. In the northern part of this extensional belt, the ages of slip on the Lewis thrust and Flathead fault indicate that crustal shortening ended in early Eocene time (ca. 55 Ma), whereas extensional faulting iiutiated in middle Eocene time (ca. 48 Ma). Tectonism and sedimentation in the BCishenehn basin continued until late Oligocene or early Miocene time (ca. 25- 21 Ma). Data related to the timing of the Medicine Butte thrust and Acocks- 59

Almy normal fault system indicate that the latest phase of thrust-fold shortening was late early Eocene (late Wasatchian; ca. 55-51 Ma) in age, and that onset of extension was in early middle Eocene time (Bridgerian; ca. 49-48 Ma). Synextensional middle Eocene-early Miocene deposits in the Kishenehn basin and Fowkes half graben are not unique in the Cordilleran fold and thrust belt. Similar basin-fill deposits are foimd in several basins in southwest Montana and Idaho (Fields et cil., 1985;, Fritz and Harrison, 1985, Hanneman and Wideman, 1991, M'Gonigle and Dabymple, 1993; Janecke, 1992, 1994), southwest Wyoming and northeast-central Utah (Oriel and Tracey, 1970; Lamerson, 1982; Royse, 1983), and, central Utah (Standlee, 1982; Royse, 1983; Riess, 1985; Constenius, 1995). However, it has recently been proposed that rocks of the Renova Formation foimd in many of the half grabens in southwest Montana are unrelated to extensional basin formation, whereas other half graber\s in Idaho and western Montana are part of a -100 km-wide, north to north-northwest trending rift superposed on the thrust- belt (Fritz and Sears, 1993; Janecke, 1994). The Renova Formation has been characterized by Janecke (1994) as fine-grained, thin, and laterally continuous, floodbasin and lacustrine deposits that formed in a broad, quiescent basin. However, these criteria alone cannot be used to discriminate between tectonic and non-tectonic deposits because; (1) Modem and ancient synextensional deposits are commonly fine-grained, especially in laciistrine-flood basin settings (for example. Lake Tanganyika, Africa; Neogene Great Salt Lake basin, Utah; late Paleogene Kishenehn Formation, Montana; Triassic and Jurassic Newark 60

Supergroup, eastern North America) (Cohen, 1990; Bortz et al., 1985; Constenius, 1981, 1989; Olsen, 1990). Ironically, the Sixmile Creek Formation, a synextensional unit overlying the Renova Formation, is lithologically so much like the Renova Formation that differentiating these units can be difficult (Hanneman and Wideman, 1991). Additionally, the fine-grained stratigraphic criteria has been inappropriately applied because the Renova Formation and its equivalent the Kishenehn Formation are found in three half grabens in the hypothesized rift, Kishenehn, South Fork, and Big Hole basins (Fields et al., 1985; Constenius, 1988, 1989). (2) The thickness of the Renova Formation measured at the surface and in wells indicate that it can be quite thick (for example, -1180 m Deer Lodge basin, ~1060+ m Jefferson Valley, -1800-3000 m Townsend Valley, and -500-2450 m Big Hole basin) (Mertie et al., 1951; Kuenzi and Fields, 1971; Fields et al., 1985; Constenius, 1988). (3) Modem and ancient lake and floodbasin deposits are known to bridge across accommodation zones and horsts in extensional settings (for example. East African Rift lakes; Qxiatemary-Neogene Lake Bonneville and Great Salt Lake areas) (Rosendahl et al., 1986; Lambiase, 1990; R.A. Johnson, pers. commim., 1995). Criteria supporting a synextensional origin of the Renova Formation are: stratigraphic thickening toward basin-boimding faults, inordinate thickness, and intercalated course-clastic deposits (Rasmussen and Fields, 1983; Fields et al., 1985). The structural setting and style of many of these basins is akin to the Kishenehn basin in that basin- boimding listric normal faults have reactivated folded and thrusted rocks (Rasmussen and Fields, 1983; E.H. Johnson, written commun., 1995). 6 I

Plotting the latitudinal distribution of ages of late Paleogene basin-fill deposits with respect to youngest foreland-basin deposits and overlap assemblages, Paleogene volcanic deposits, age-ranges of metamorphic core complex extension, and ages of Neogene basin-fill deposits reveals the following time-space trends (Fig. 2.10): (1) Dating of yoimgest foreland-basin deposits in the fold and thrust belt between latitudes 40® 30'N and 490N indicates the end of contractile deformation occurred between 54 and 51 Ma. Foreland-basin sedimentation south of 40° 30'N ended in late Campanian to early Maastrichtian time (Lawton, 1985), but, deposition of overlap assemblages consisting of the North Horn, Colton, Green River and Uinta formations was continuous from latest Cretaceous to late middle Eocene (ca. 42-40 Ma) time. Stratigraphic data from basins in the Foreland indicate crustal shortening ended in early and middle Eocene (ca. 55-50 Ma) time north of 420N and late Eocene (ca. 40-35 Ma) time south of 420N (Dickinson et al., 1988). (2) Middle Eocene volcanic deposits post-dated the end of foreland basin sedimentation and preceded sjniextensional basin-fill sedimentation. Magmatism and basin-fill sedimentation were temporally discrete north of 430N (minor coeval magmatism and extension) but were coeval in late Eocene to late Oligocene time south of 430N. (3) The hiatus between contractile and extensional sedimentation where narrowly bracketed was brief, 5 Ma in north and 2 Ma in south. However, throughout much of the Cordilleran foreland, the hiatus is long (ca. 16-25 m.y.). 62

Figure 2.10. Chart showing age distribution of late Paleogene basin-fill deposits (solid gray) v^dth respect to youngest foreland basin deposits (black candy cane), volcanic deposits (black), age-ranges of metamorphic core complex extension (white rectangles with black border), and ages of Neogene basin-fill deposits (gray candy cane) in the Cordilleran foreland fold and thrust belt. Vertical axis is time labeled in Ma. Horizontal axis is latitude. Extension in fold-thrust belt (1) closely follows end of crustal shortening, (2) overlaps with formation of Cordilleran metamorphic core complexes, (3) mostly post-dates regional magmatism north of 40O30'N (Archean- Proterozoic crustal boundary), (4) is coeval with magmatism south of 40O 30' N, and (5) is temporally distinct from Basin and Range extension. Regional Hemingfordian-age imconformity marked by ca. 20-16 Ma hiatus (Rasmussen, 1973, Fields et al., 1985). Sources of stratigraphic-age data: Russel (1950,1968), Oriel and Tracey (1970), Qague (1974), Nelson (1973,1974, and 1979), Young (1976), Dorr et al., (1977), Miller (1980), Marvin et al., (1980), Standlee (1982), Wiltschko and Dorr (1983), Maeser et al., (1983), Crittenden and Sorensen (1985), Fields et al., (1985), Chadwick (1985), Lawton (1985), Van Horn and Crittenden (1987), Hintze (1988), Hartman (1989), Rasmussen (1989), Witkind and Marvin (1989), Bryant et al., (1989), Constenius et al., (1989), Nichols and Bryant (1990), Bryant (1990), Fox (1990), Hanneman and Wideman (1991), Lange and Zehner (1992), Janecke (1992), Janecke and Snee (1993), Lillegraven (1993). Sources of data for metamorphic core complexes: Coney (1980), Saltzer and Hodges (1988), Armstrong and Ward (1991), Lee and Sutter (1991), Harms and Price (1992), Burchfiel et al., (1992), Axen et al., (1993), Hodges and Applegate (1993), Mueller and Snoke (1993), Wells and Snee (1993), TIME -SPACE PATTERN OF TECTONISM

CRETACEOUS PALEOCENE Thrusting - foreland basin sedimentation EOCENE ML

Sevier belt collapse OLIGOCENE Metamorphic core complexes

MIOCENE Basin & Range extension PLIOCENE o o 49 47 45 43 41 39 LATITUDE PROTEROZOIC CRUST 64

(4) Extension in the metamorphic core complexes in the hinterland of the Cordillera partly overlapped with sjmextensional basin-fill sedimentation in the foreland north of 430N. Extension in the foreland and hinterland was concnrrent south of 430N. (5) Late Paleogene synextensional sedimentation records an episode of tectonism discrete from older foreland basin sedimentation and later Basin and Range deformation. Basin-fill sedimentation spanned from middle and late Eocene to early Miocene (ca. 49-21 Ma) north of latitude 40° 30' north. South of 40° 30'N, basin-fill sedimentation starts in late Eocene time (ca. 39 Ma), about 10 m.y. later than the start of synextensional sedimentation to the north, and ends in the early Miocene. (6) A regional Hemingfordian-age unconformity (Rasmussen, 1973; Fields et al., 1985) separates late Paleogene synextensional deposits from younger, middle Miocene to Recent (17 to 0 Ma), deposits that were the product of Basin and Range extension and magmatism. The hiatus associated with this imconformity is about 5 million years.

2.6 LATE CRETACEOUS -PALEOGENE MAGMATISM - A RECORD OF CHANGING PLATE REGIMES

Studies of the Cordillera in the southwestern U.S. and northern Mexico have used the concept of a "magmatic sweep", that is the migration of arc-magmatism spatially through time, to draw insights into the timing and mode of Cenozoic tectonism. This concept was originally applied by Coney 65

and Reynolds (1977) who documented systematic Late Cretaceous and Tertiary shifts in magmatism in the southern Rocky Moimtains and proposed that the ~1000-km-eastward migration of the magmatic arc was related to Late Cretaceous and early Tertiary crustal shortening concomitant with flattening of the subducted oceanic slab beneath North America. Late Cretaceous- Eocene crustal shortening in the Cordillera has also been linked to: (1) high relative plate convergence rates (greater than 70-150 km/m.y.), (2) subduction of abnormally buoyant lithosphere, (3) subduction of very young oceanic lithosphere, and/or (4) increased rates of absolute motion of the overlj^g plate toward the trench (Cross and Pilger, 1982; Miller et al., 1992 and references therein; Ward, 1995). The validity of the dipping-slab hypothesis (Coney and Reynolds, 1977) has been challenged by Ward (1991,1995) who also demonstrated the hazards of generalizing from "magmatic sweep" diagrams. Middle Tertiary silidc volcanism and metamorphic-core-complex- related extensional deformation in western North America has been correlated with westward migration of arc-magmatism (for example, Dickinson, 1991; Steme and Constenius, 1997). This pivotal change from contractile to extensional deformation may have resulted from one, or a combination of the following factors: (1) slowing of relative Pacific-North America plate convergence rates, (2) age of the subducted plate, (3) a change to subduction at steeper angles beginning in middle Eocene time, and (4) subduction of a linear aseismic ridge (Wernicke, 1992 and references dted therein; Dickinson et al., 1988; Engebretson et al., 1985; Ward, 1995). Patterns of migrating arc-magmatism are viewed here as a tectonic signal of rates of plate convergence and state of the subducted slab, and are tised to delineate the timing of contractile versus extensional regimes in the northern U.S. Cordillera. Published isotopic ages of late Cretaceoxis through Eocene igneous rocks, age-of-deformation data, and stratigraphic-age data were compiled for the northern U.S. Cordilleran foreland on a line of projection orthogonal to tiie orogenic belt (N6OOE) (Figs. 2.11 and 2.12). Deformational-age and stratigraphic-age data were included to compare tectonic regimes predicted from the magmatic-sweep diagram and to place bounds on interpretations. Relationships interpreted from Figure 2.11 are: (1) accelerated plate convergence rates and/or flattening of the subducted slab were concomitant with -600 km eastward progression of magmatism and crustal shortening during Late Cretaceous to early Eocene (ca. 72-54 Ma) time; (2) in early-middle Eocene time (ca. 53-51 Ma) markedly reduced plate convergence rates and/or slab-steepening resulted in westward migration of the subducted slab (rollback) and inception of widespread magmatism; and (3) middle Eocene (ca. 49-48 Ma) initiation of normal faulting and basin-fill sedimentation in the Cordilleran thrust belt was related to westward passage of the subducted slab from beneath the Great Plains to a position closer to the Pacific coast (Fig. 2.12). Inspection of Figure 2.11 reveals that the initiation of magmatism is about 5 -6 million years yoimger in central Idaho than in central Montana, 600 km to the east. The rate of slab rollback based on this age-distance comparison is about 100 km/m.y. This interpretation is bracketed by dating phases of contractile and extensional deformation in the thrust belt and cessation of foreland basin sedimentation (Wasatch Figxire 2.11. Magmatic sweep and initiation of extensional basin formation in northern Rocky Mountains. Ages of igneous activity, foreland basin deposits of Wasatch Formation, and crustal contraction and extension plotted with respect to distance (modified from Lipman, 1992) from western margin of North American plate. Extensional basin formation in Cordillera begins ca. 49-48 Ma with westward passage of late Paleogene regional magmatism that tracks rollback of subducted slab. Dataset comprises predominantly published ^Ar/39Ar and K/Ar dates (McDowell,1971; Marvin et al., 1973; Marvin et al., 1980; Oiadwick, 1980; Armstrong et al., 1982; Marvin et al., 1982; Staatz, 1983; Marvin et al., 1983; Qiadwick, 1985; Qiesley, 1986; Whipple et al., 1987; Meen and Eggler, 1987; Harlan et al., 1988; Bellon, 1989; Harlan et al., 1992; Jolly and Sheriff, 1992; Harms and Price, 1992; Janecke and Snee, 1993, and Hodges and Applegate, 1993). K/ Ar ages include only those with less than fifteen percent analytical error. Data points projected onto the line trending N650E shown in Figure 2.12. The open circle data point is the potassixmi-argon date of 59.6 +/- 1.6 Ma from a dike that intruded across Steinbach thrust zone - a critical date regarding historical reconstruction of tectonic events. Southivest IDAHO EASTERN LIMIT OF THRUSTING Northeast BOULDER HIGHWOOD MOUNTAINS BLACK !: CRAZY BEARPAW HILLS MOUNTAINS MOUNTAINS

- |S TIME OF THRUSTING (-7J.S7 Ma) £;uj (HOFFMAN ET AU 1976^ oB 5S YOUNGEST DISPLACEMENT z.< UN LEWIS THRUST (-65-55 YOUNGEST eSc (COVEY ET AU 19941 SEDIMENTARY ROCKS Bt-i DERIVED FROM THRUST BELT O" (WASATCH FM^ -57-54 u)X FOWKES h WOODRUFF — NI-WrORT FAULT HAIF GRABENS (TIGER FMI " KISHENEHN If BASIN

AGES OF UNDEFORMED z INTRUSIVE ROCKS IN THRUST BELT* EASr-CENTRAL IDAHO o ui (HARLAN ET AU 1900) CO HALF GRABENS zZH^ I/) s s HORSE TRAIRIE-MEOICINE LODGE BASINS X X I "tlJ Hm J- X -i- 300 400 500 600 700 800 900 1000 1100 DISTANCE FROM TRENCH (km) 69

Figure 2.12. Cordilleran tectonic elements related to late Paleogene cnistal extension. Collapse of Cordilleran orogen was widespread and development of metamorphic core complexes and low-angle detachment faults in hinterland was partly synchronous with listric normal faulting in foreland fold and thrust belt. Extensional collapse of fold and thrust belt manifested as network of late Paleogene half graben superposed on this structural belt. Sources of data for extensional structures dted in Figiu*e 2.11. Sources of data for magmatic belts Qiadwick (1985), Armstrong and Ward (1991), Christiansen and Yeats (1992). Contours of subducted slab positions (S=2; aseismic slab that may affect tectonics of continental lithosphere) from Severinghaus and Atwater (1990). Bearpaw Mountains (BPW). Crazy Mountains (CZM). Highwood Moimtains (HGW). Pioneer metamorphic core complex (PNR). Ruby Moimtains-East Humboldt-Wood Hills metamorphic core complex (RB-EH-WH), 70

'rv CORDILLERAN FORELAND FOLD & THRUST BELT

rv KtSHENEHN oc:^cp^'°^ BASIN !«. ^ uSA. (48-21 Ma) PRIEST RIVER

(SS.42 HR,,"^T HGW

M 'R O"*' •^OO' %p E-^ST-CrvTB V '«3- " >/ '^'NTRalHiirr IDaho & O'C

^'ON-RAFT RIVED "{SCAR RA 'A « FOWKES-WOODRUFF "3-37 Man HALF GRABENS RB-EHCW# P* (49-27 Ma) MAJ

SIERRANMAGMATIC-WAS BELTATCH V • RANGE (37-21 Ma)\ \ >• *

"^-10 Ml POSITION r • -SOMaPJ^mON OF SUBDU

1 -\

"ORTH

500 km

CORDILLERAN FORELAND FOLD O AND THRUST BELT FORELAND PROVINCE

LATE PALEOCENE MAGMATIC BELTS \LATE PALEOCENE HALF GRABENS AND ' GRABENS IN THE THRUST BELT -C oMETAMORPHIC CORE COMPLEX 7 1

Formation). Independent dating of deformational phases agrees with tectonic regimes predicted from the magmatic sweep diagram; that is, eastward migration of arc-magmatism is equated with crustal shortening, and westward migration of arc-magmatism correlates with late Paleogene magmatism and subsequent extension. Models of North America-Pacific basin plate convergence reported by Engebretson and others (1985) and Cole (1990) show a dramatic change from rates as high as -120-150 km/m.y. and near orthogonal convergence in Late Cretaceous to early Eocene time, to rates as low as -50-86 km/m.y. and convergence with an oblique component in late Eocene to early Miocene time. Extensional basin formation in both the foreland and hinterland of the Cordillera is the structural response of a crust preconditioned by contractile deformation which has been left unsupported by lower rates and changes in orientation of plate convergence concomitant with slab rollback. Passage of the subducted slab from beneath the Great Plains to the Cordillera at ca. 49-48 Ma, signals widespread breakup of the Cordilleran orogen which collapsed xmder its own weight and spread to the west. Magmatism preceded extensional basin formation by -1-3 m.y. in the Horse Prairie-Medidne Lodge and East Central Idaho half grabens (M'Gonigle and Dalrymple, 1993; Janecke, 1992 ,1994), whereas, there are no middle Eocene volcanic fields within a -100 km radius of the Kishenehn basin and Fowkes half graben. These observations preclude a genetic link between local magmatism and extension (10-100 km radius; Axen et al., 1993) but are consistent with thermal weakening of the crust and lithosphere and/or changes in stress transmitted from the subducted slab. 72

2.7 EXTENSIONAL MECHANISMS

Extension of supracrustal rocks in the Cordilleran foreland fold and thrust belt occurred in areas that had been thickened during crustal shortening and subsequently failed along weaknesses imparted by thrust-fold deformation due to their high gravitational potential. Extensional collapse preferentially reactivated pre-existing thrust surfaces and was superposed on large-relief thrust-fold structural features such as structural culminations, duplex zones, thrust ramps, and associated fault- bend folds. The principal factors that increased the gravitational potential and facilitated late Paleogene normal faulting are: (1) inherited structural weaknesses in the rock column such as pre-existing faults and physically incompetent rock units , (2) large structural relief created by stacking of thrust-fold structures, (3) the structural framework of the hanging wall (loi^g^ west-dipping panels of strata), and (4) the 3-6° west-dip of the sole fault. Irtspection of Figures 2.4, 2.6, 2.8 and 2.9 reveals that late Paleogene normal faulting is superposed on major thnast-fold structures with large vertical dimensions. The ramp-geometry of the Lewis and Medicine Butte thrust surfaces combined with the formation of footwall duplex zones created structural relief of about 4 and 7 kilometers, respectively. The critical aspect of these structures with respect to crustal instability, however, is that they formed ~25-km-long, ~ 150 west-dipping panels of anisotropic strata weakened by detachments. The enhanced structural relief and resultant increase in the potential energy of Lewis and Medicine Butte hanging wall 13 masses drove backsliding on the detachment siufaces. Additionally, certain stratigraphic imits are ideal slip surfaces for accommodating both thrusting and low-angle normal faulting. The Medicine Butte fault is mainly an evaporite detachment, whereas ttie Lewis detachment juxtaposes argillite, siltite, quartzite, and dolomite on shale, sandstone, and limestone. In the case of the Medicine Butte thrust, the Preuss evaporite unit not only is the preferred horizon for the sole fault, but it also may be the mechanically weakest part of the entire stratigraphic colimm. Recent literature regarding extension and low-angle normal faulting in active orogens such as the Himalayas-Tibet and Andes provide Neogene analogs of the role of gravitationally driven extension in high relief areas (Burchfiel et al., 1993; Merder et al., 1992; Sebrier et al,, 1985). The ancient Cordillera differs from these modem high mountain chains in that there is presently no evidence of extension coeval with Late Cretaceous to early Eocene crustal shortening in the foreland fold and thrust belt. Studies of the age of normal faults in the Idaho-Wyoming-Utah thrust belt have concluded with one possible exception, that all the normal faults formed subsequent to thrust-fold deformation (Armstrong and Oriel, 1965; Wiltschko and Dorr, 1983). Furthermore, the detachments which accommodated the collapse of the orogen were deep in the subsurface rather than exposed at the surface in topographically high areas. The late Cretaceous to early Eocene paleotopographic relief of the Lewis thrust sheet west of the Flathead fault is difficult to assess, but, the inferred presence of highly erodible Cretaceous rocks in the hanging wall of the Lewis thrust implies low local relief. Correlation of Paleocene and early Eocene stratigraphic units across the 74

Medicine Butte thrust indicate that it had comparatively little paleotopographic relief. On a regional scale, however, the Lewis thrust sheet and Wasatch culmination were major contributors of coarse-clastic sediments to the foreland basin which implies that they had significant paleotopographic relief (Heam et al., 1964; DeCelles, 1994).

2.8 STRUCTURAL GENESIS

Comparison of the northwest Montana and southwest Wyoming- northeast Utah extensional systems reveals similarities in the style and timing of tectonism and yields insights into local and regional processes that shaped the Cordillera. These two areas, though widely spaced, record an episode of regional gravitational collapse of the Cordilleran foreland fold and thrust belt that initiated in the early middle Eocene (ca. 49-48 Ma) and ended in early Miocene time (ca. 20 Ma), Concurrently, in the Cordilleran hinterland, deep-seated crustal extension and magmatism reshaped an even thicker part of the gravitationally unstable orogenic wedge (Coney and Harms, 1984; Coney, 1987; Harms and Price, 1992). Large-scale crustal extension in hinterland areas has been attributed to gravitational spreading of thickened, structurally preconditioned, and thermally weakened crust triggered by changes in the regional stress regime (for example. Harms and Coney, 1984; Axen et al., 1993). The correspondence of extension and westward sweep of late Paleogene magmatism presented here suggest collapse of the Cordilleran orogen was related to slowing of plate convergence rates and steepening of the subducting slab (Figs. 2.11 and 2.12) (Lipman, 1992; 75

Severing^us and Atwater, 1990). Linkage of late Paleogene extensional processes that simultaneously deformed both the hinterland and foreland of the Cordillera may be conceptualized using the Coulomb critical taper hypothesis (Davis et al., 1983). Contemporary thinking regarding contractile deformation, propagation, and structural thickening of thrust belts and accretionary wedges has effectively relied on the Coulomb critical taper model, and conversely, the extensional wedge concept has been used to explain regional continental extension (Xiao et al-, 1991). Extensional wedges are defined by: (1) a gently dipping basal detachment down which the tapered hanging wall slides, (2) overlying the basal detachment a related system of normal faults cuts the wedge, and (3) the footwall beneath ttie basal detachment is relatively undeformed (Wernicke, 1981; Xiao et al., 1991). Normal faxilt systems superposed on the Cordilleran foreland fold and thrust belt satisfy these criteria as shown in Figures 2.6 and 2.9. Extensional wedges are the inverse of compressional wedges, but both share the relationship that ttiey are critically tapered when on the verge of failure throughout the wedge. An extensional wedge whose taper exceeds a certain critical taper deforms by internal normal faulting and reduces its taper imtil it is critical. If the taper of the extensional wedge is narrower it can slide without internal deformation down the inclined basal detachment (Xiao et al., 1991). In late Paleogene time, the dramatic reduction in east-west horizontal compressive stress at ca. 50 Ma left the Cordilleran orogenic wedge largely imsupported to the west. Consequently, the orogen switched from a state of failure imder shortening to one dominated by extension and it 76 gravitationally collapsed and horizontally spread to the west iintil an eq\jilibrium was established at ca. 20 Ma. The Lewis thrust salient and the Wasatch c\ilmination were predisposed to large-sccde normal faulting and development of late Paleogene half grabens because these were areas of heightened structiaral and paleotopographic relief that required considerable extension to achieve critical taper.

2.9 CONCLUSIONS (1) Development of late Paleogene extensional basins in the Cordilleran foreland fold and thrust belt overlapped or was coeval v^th formation of the metamorphic core-complexes, low-angle detachment faulting , and regional- scale magmatism. Extensional structures in the foreland like the Kishenehn basin were not formed as discrete and isolated entities, but are a manifestation of late Paleogene gravitational collapse and breakup of the entire Cordillera. The Cordilleran foreland fold and thrust belt failed as an extensional wedge governed by Coulomb critical taper criteria.

(2) The beginning of extension in the Cordilleran foreland fold and thrust belt closely followed the termination of contractile deformation after a brief hiatus. Even though rollback of the subducted Pacific plate and concomitant middle Eocene magmatism initiated in the northern Great Plains at ca. 53-51 Ma, the period of extensional basin formation and collapse of the Cordillera commenced at ca. 49-48 Ma.

(3) In the Cordilleran foreland fold and thrust belt, the late Paleogene episode 77 of extensional collapse and the Basin and Range extensional event are two distinct episodes of cnistal extension whose sedimentary successions are separated by a major angular unconformity ( 3-15 Ma hiatus). Many of the present topographic basins, which were constructed largely during Basin and Range-Recent extension, have variously buried, broken-up and/or reactivated the earlier late Paleogene half grabens and grabens. 78

CHAPTER 3 STRUCTURE OF THE DEER CREEK DETACHMENT FAULT SYSTEM AND EVOLUTION OF THE COTTONWOOD METAMORPHIC CORE COMPLEX, WASATCH MOUNTAINS, UTAH

3.1 ABSTRACT The Deer Creek fault is an extensionally reactivated sole-thnist linked to a network of former thrust surfaces that collectively accommodated late Paleogene extensional failure and at least 5-7 km of west-southwest transport of the C-N allochthon. Unique "bowl-shaped" half-grabens located in the southern part of the Charleston-Nebo (C-N) allochthon are superposed on detachments inherited from a triangle zone. Preserved in these half-grabens is a long record of C-N extension. Age-dating and paleontologic data suggest that the sole fault of the allochthon was extensioiudly reactivated in late Eocene time (ca. 38-40 Ma) and displacements continued throu^ middle Miocene time (ca. 20 Ma). Evidence of this early extensional episode along the north margin of the C-N allochthon is seen in extensional growth strata of the late Eocene- Oligocene Tibbie Formation, Oligocene dikes cut by detachment faults, an immense rollover structure (Box Elder Peak anticline), and rapid cooling of deep-seated, tectonically denuded stocks in the Wasatch igneous belt (WIB). Tephras from the lower and upper parts of the Tibbie Formation yield ^Ar/39Ar ages of 36.4 +/- 0.2 Ma and 28.5 +/- 0.2 Ma, respectively. Structural and paleobarometric data indicate the basal Tibbie imconformity at the west end of the half-graben has been down-dropped in excess of 7 km relative to the uplifted and denuded Deer Creek footwall terrain. 79

Slip on the Deer Creek detachment system was synchronous with formation of metamorphic core-complexes and low-angle detachment faulting in western Utah and Nevada. The correspondence of structural elements, magmatism, synextensional sedimentation and timing of the Deer Creek fault system with core-complex-related detachment faults suggests that the Deer Creek system is a core-complex-related fault system linked to the collapse and westward transport of the C-N allochthon and flexural-isostatic rebound of the Deer Creek footwall. This system is introduced here as the Cottonwood metamorphic core complex. Interpretation of mineral assemblages and fluid inclusion data in the WIB and Comer Creek tectonite (Parry and Bruhn, 1986; John, 1989) shows that the Cottonwood complex is an oblique cnistal profile with paleodepths of ~ 11 km in the west, progressively shallowing to less than 1 km in the east. We hypothesize that this differential and large magnitude of flexural-isostatic exhumation and WIB crusted melts formed in response to reboimd of a Cheyenne belt crustal welt that was tectonically denuded by the Deer Creek detachment.

3^ INTRODUCTION The Charleston-Nebo (C-N) allochthon. Deer Creek detachment fault, Wasatch igneous belt (WIB), Cottonwood metamorphic core complex, and related tectonic elements of tiie central and southern Wasatch Mountains are superposed on the crustal boundary between Archean-early Proterozoic (ca. 2.4-2.7 Ga) rocks of the Wyoming province to the north and Proterozoic (ca. 1.6-1.8 Ga) rocks of the Yavapai province to the south (Fig. 3,1; Bryant, 1988; Houston et al,, 1989; Presnell, 1997). This crustal suture, named the Cheyenne 80 belt (Houston et al., 1979), developed as a zone of large-scale thrusting followed by strike-slip faulting, during Proterozoic accretion of oceanic rocks of the Yavapai Province to cratonic rocks of the Wyoming Province. The resultant structure is an E-W oriented boundary zone that extends from eastern Wyoming to northeastern Nevada (Roberts, 1965; Wright and Sboke, 1993; Stone, 1993). Recurrent crustal movements on the Qieyerme belt are recorded in the tectonostratigraphic record of the Precambrian-Paleozoic rock column, with the most prominent accumulation of faulted strata constituting the 6-10 km thick Proterozoic Big Cottonwood-Uinta Mountains Group. The jimction of the Cheyenne belt with normal faults related to late Paleogene collapse of the Sevier belt and Neogene Basin and Range extension (Wasatch fault) is marked by considerable crusted extension and magmatism. Igneous intrusior\s of the late Paleogene WIB are aligned v^dth the trend of the Cheyenne belt (locally called the Uinta-Cortez axis) as are the axes of the Cottonwood arch and Uinta Mountains uplift (Presnell, 1997; Vogel et al., 1997). The most recent tectonomagmatic expression of this crustal flaw is the Cottonwood metamorphic core complex (new ncime), a feature comprised of tj^ical core complex elements: detachment faults and related fault-rocks, synextensional plutons, marble tectonites, extensional growth strata, and flexural-isostatic exhumation of formerly deep-seated rocks (Coney, 1980). In this study we investigate the causative mechanisms that lead to construction and later structural collapse of the C-N aUochthon and development of the Cottonwood metamorphic core complex. We conclude that: 1) Sevier belt fold-thrust structures provided the structural template for linked contractile-extensional faults systems, 2) the basal thrust of the C-N Figure 3.1. Geologic index map of major tectonic elements in north-central Utah showing location of geologic cross sections, structures, seismic profile and wells (circled numbers and letters) and gravity profile. Undifferentiated Paleozoic and Mesozoic rocks of Charleston-Nebo allochthon shaded grey. Position of Nebo thrust imbricate (Park Gty Formation hanging wall cutoff) formerly considered leading-edge of C-N thrust system by Baker (1972) and Witkind (1987) shown. Inferred position of the Qieyenne belt show (candy cane line). Abbreviated names: D, fensters of Deer Creek thrust; H, Hunt Indianola Federal 24-1 & Union Brown's Peak Federal l-G-24 well; P, Payson fenster; R; Red Rock Meadow; S, Sohio Indianola 1 well; T, Thistle, Utah; U, Union Federal J-9 well, and, W, White Lake Hills thrust. See Figure 2 for rock tmit abbreviations. Map adapted from Hintze (1980). 82

Keetle Wasatch / volcanic fault

Salt Lake City • p V ^ Cottonwoo

Cottonwood I 111 ocw metamorphic| enjtfaulf 4-4(f30'N core complex' (Figure 7) Triangle zone Tibbie half-graben Provo Little Diamond Creek half-graben CN Uinta basin-mountain boundary fault (blind thrust) Santaquin Meadovrs & Nebo thrust imbricate Payson Lakes huf-grabens Charleston- Nebo © thrust f N Little Clear Creek 20 km half-graben 83 allochthon was extensionally reactivated in late Paleogene-early Neogene time (ca. 40-20 Ma) and accommodated at least 5-7 km of west-southwest tectoruc transport, 3) the signature of late Paleogene collapse is seen in detachments faults, half-grabens, rollover anticlines and s3aiextensional igneous intrusive bodies of the Wasatch igneous belt, 4) WIB magmas were the products of crustal melts derived from melting of lower crustal rocks that had been tectonically tiiickened and depressed, 5) the same crustal welt that was the soxirce of WIB magmas is also the cause of flexural-isostatic footwall uplift on the Wasatch-Deer Creek fault system, and 6) that a relict of this once large crustal welt expresses itself in Bouguer gravity data. In support of tiiese hjrpotheses and to document the timing and style of tectonics related to the collapse of the C-N allochthon and the evolution of the Cottonwood metamorphic core complex, we present geologic cross sections based on revised mapping and interpretations, stratigraphic and structural-kinematic data, a seismic reflection profile, geochronologic results, and a Bouguer gravity model. Our data and analyses are presented in three sections. The first examines the mode of contraction sustained by the C-N thrust during the Sevier orogeny that created an inclined structural fabric prone to collapse. Second, we docmnent the timing and structural signatture of late Paleogene collapse of the C-N allochthon related to extensional reactivation of the sole fault (named Deer Creek detachment fault, where it crops out along the uplifted and deformed northern margin of the thrust sheet). Third, this paper investigates extensior\al structures of the Cottonwood core complex and the history of tectoruc denudation of the Deer Creek-Wasatch footwall, and proposes mechanisms to accoimt for breaching 84

of the C-N sole fault along with exhumation of deep-seated fault-rocks and sjmextensional plutons.

3.3 FOLD-THRUST STRUCTURE: STRUCTURAL TEMPLATE FOR COLLAPSE The C-N allochthon is an immense leading-edge structural element of the Sevier belt located in central Utah (Tooker, 1983; Gallager, 1985). Dimensionally, the eroded remnant of the allochthon east of the Wasatch fault forms a 3700 square kilometer thrust salient. The stratigraphic thickness of Proterozoic to Lower Qretaceous rocks in the hanging wall plates is as much as 16.5 km (Baker, 1959; Reiss, 1985), and more than 100 km of eastward transport is estimated from palinspastic restoration (Figs. 3,2 and 3,3). The C- N allochthon consists of two large thrust sheets, the Qiarleston and the Nebo, that are the upper horses of a crustal-scale antiformal duplex referred to here as the Santaquin culmination. Collectively, these tiirusts have inverted rocks of the Pennsylvanian-Permian Oquirrh basin and transported them approximately 80 km eastward (Fig. 3,3). The resultant allochthon juxtaposes a 15-16.5 km thick miogeoclinal rock colimm with a 4.6 km thick cratonic sequence. Similarly, upper Mississippian-Permian strata of the C-N allochthon range from 5-11 km thick, whereas the same age rocks of the Cottonwood arch attain of thickness of only 1.2-3,6 km (Fig. 3,2; Baker, 1964), The inception of fault displacement on the C-N thrust system is difficult to gauge but may be as old as Upper Jurassic (ca. 155 Ma) based on dating of fault gouge and provenance studies (Presnell and Parry, 1995; Currie, 1998). Schwans (1995) dtes a Barremian age (ca. 120 Ma) as the time of initial displacement on the C-N tiirust. The first stratigraphic indication of C-N 85

Figure 3.2. Balanced geologic cross section CN-CN' of the Qiarleston-Nebo allochthon showing the Santaquin culmination, an antiformal duplex comprised of three thrust sheets, and retrograde motion on east-dipping backthrusts and other thrusts that formerly accommodated crustal shortening in a large-scale triangle zone. Thrust flats developed in Manning Canyon Shale, Park Qty Formation (base and Woodside Shale at top), and Arapien Shale that accommodated considerable horizontal shortening and later allothchon collapse shaded grey. Late Campanian-early Eocene synorogenic strata related to final phase of thrusting shaded dark grey. Geologic maps used in cross section: Pinnel (1972), Rawson (1957), Rimyon (1977), Witkind and Page (1983), Witkind and Weiss (1991), Young (1976). Rock imit abbreviations (oldest-youngest): pCxa, Archean-Early Proterozoic crystalline basement; pCx, Proterozoic crystalline basement (Little Willow Formation or Santaquin Complex); pCs, Proterozoic Big Cottonwood Group; C, Cambrian rocks undifferentiated, CDu, Cambrian-Devonian rocks undifferentiated; Mu, Mississippian rocks imdifferentiated; Mmc, Mississippian Manning Canyon Shale; PPoql & PPoqu, lower & upper parts of Permsylvanian-Permian Oquirrh Group; Pdc, Permian Diamond Creek Sandstone; Ppc, Permian Park Qty Formation; Tru, Triassic undifferentiated; Trwt, Triassic Woodside Shale & Thaynes Formation; Tra, Triassic Ankareh Formation; Jn, Jurassic Navajo Sandstone; Jtc, Jurassic Twin Creek Limestone; Ja, Jurassic Arapien Shale; Jtg, Jurassic Twist Gulch Formation; Jsc, Kcm, Cretaceous Cedar Mountain Formation; Ki, Cretaceous Indianola Group; Kd, Cretaceous Dakota Sandstone; Kmc, Cretaceous Mancos Shale Group, BCsp, Cretaceous Star Point Sandstone; Kbh, Cretaceous Blackhawk Formation-South Flat Formation; Kc, 86

Cretaceous Castlegate Sandstone & Price River Formation; KTnh, Maastrichtian-early Tertijiry Nortii Horn Formation; T£, early Eocene Flagstaff Limestone; Tgr/Tu, middle Eocene Green River and Uinta formations; Tm, late Eocene-middle Miocene Moroni Formation; Tv, late Eocene-Oligocene igneous rocks of Keetley volcanic field; Ti, late Eocene-Oligocene plutonic rocks of Wasatch igneous belt, and, Ts, Miocene-Recent basin-fill. CN West Santaquin culmination CN' Vast Wamtch Fork Witmich fhuit MounMiu "ilf-graben Clear Creek Hollow Emery uplift msu.,„n J"'f'l^oten anticline DffrCm* UiTuat '^"'flinuloflriannlezow Dairy Fork jhult Plateau Sohio Indianola #7 IhttUt / Cn trm>/Kt,d ?.OtmNE) Xnjo/i/i finita

Uinla BMB thrust

Ihvl ZIf-c •"-5 km Pile «• --10km

km

00 -vj 88

Figure 3.3. Structioral evolution of the Charleston-Nebo allochthon illustrating development of triangle zones and east-dipping passive roof thrusts developed in mechanically weak imits. Detachments that were subsequently reactivated as normal faults during late Paleogene-early Miocene collapse of the allochthon. Ehiring early phases of thrusting a triangle zone develops in evaporites of the Jurassic Arapien Shale. Growth of the Santaquin culmination began when thrust-tip displacement stalled. Henceforth, the allochthon sustained intraplate contraction as the leading- edge Arapien triangle zone was breached and abandoned. The "intraplate" triangle zone developed passive roof thrusts in the Triassic Woodside Shale, Permian Park Qty, and Mississippian Manning Canyon Shale above an intercutaneous wedge of the Santaquin culmination (Qiarleston and Uinta BMB thrust sheets). Large eastward displacement of the wedge (ca. 50 km) results in stratigraphic attenuation across segments of the reactivated Nebo thrust, because in these places the sense of hanging wall-footwall thrust displacment has been reversed. Compare position of Pennsylvanian-Permian reference point "r" with respect to Proterozoic-lower Paleozoic rocks of the Nebo thrust hanging wall. Rock imits: (l)Proterozoic crystalline basement (black), (2) Cambrian-Mississippian (dark gray), (3) upper Mississippian- Permian (white), and, (4) Triassic-Jurassic (gray). 89

UINTA BMB THRUST East Santaquin culmination 25 r (-82-40 Ma, 20 km) West

25 km

CHARLESTON THRUST - 2 (-87 - 84 Ma, 35 km)

CHARLESTON THRUST- 1 (-90-S7Mtt, 15 km)

NEBO THRUST (-155(7) /120(?) - 90 Ma, 35 km)

Charleston thrust sheet Nebo thrust sheet Uinta BMB sheet Emery uplift 90 thrusting was considered by Jefferson (1982) and Lawton (1985) to be the Albian basal conglomerate of the Indianola Group (ca. 110). Mitra (1997) cautioned that the source of the majority of dasts in the lower Indianola Group were more likely derived from thrust sheets west of the Qiarleston- Nebo thrust. Stratal growth in coarse clastic rocks of the Lower Cretaceous Indianola Group in the hanging walls of the Qiarleston and Nebo thrusts near Thistle, Utah, and fault truncation of Campanian rocks in the footwall of the Charleston thrust, indicates that these faults were active during Coniadan through early Campanian time (ca. 90-84 Ma) (Fig. 3.3; Lawton, 1985; Bryant and Nichols, 1988; Mitra, 1997). Thrusts of the C-N system are erosionally truncated and buried by Campanian strata of the Blackhawk-South Flat, Castlegate-Price River, and Currant Creek Formations (Figs. 3.2, 3.3, and 3.4; Walton, 1944; Garvin, 1969; Isby and Picard, 1985; Bryant and Nichols, 1988; Haldar, 1997; Mitra, 1997). This brackets the time of last movement on these thrusts at about 80 Ma. Interpretation of reflection seismic profiles, down-structure views of the outcrop belt and construction of a balanced geologic cross section suggest that the Charleston and Nebo tiirust faults cut through the rock colimm as follows. The thrusts ramp upsection through mechanically competent formations in four major steps (Figs. 3.2, 3.3 and 3.4; Bryant, 1992): 1) from Proterozoic crystalline basement and Proterozoic Big Cottonwood Formation through the Mississippian Great Blue Limestone, 2) from the base of the Peraisylvanian-Permian Oquirrh Group (sandstone and limestone ca. 4-7 km thick) through the Permian Diamond Creek Sandstone, 3) from Triassic Figure 3.4. Migrated and depth converted seismic reflection profile 12. Subsurface structure imaged on this true-scale profile incorporated in construction of bzdanced cross section CN-CN'. Correlation of formations based on synthetic seismgrams from boreholes and projection of surface geology linked to a grid of 2-D sesimic reflection profiles. Intersection points of 2-D seismic lines from data grid marked by black triangles. See Figure 3.2 for abbreviations. Seimic reflection profile courtesy Exxon Exploration Co. 92

Little Clear Creek Oil Hollow Emery uplift Little Diamond Fork anticline half-graben half-graben Hunt Inditmola #24-1 Wasatch Plateau Little Diamond (pnjected S.OhnNNE) Pork fault Rollover anticline Thistle fault Dairy Fork fimlt Sohh Indianala tl West Union 1-9 (projected 7.0 km NE) (projected S.OhnNE) East ^0

10

Little Clear Creek Oil Hollow Little Diamond Fork half-graben Emery uplift half-graben anticline Hunt tndianola #24-1 Wasatch Plateau Little Diamond Rollover anticline (projected 8.0kmNNE) Fork fimlt Thistle fimlt Dairy Fork fimlt Sotop ImUanoUt *1 West \ Vdsiu Union 1-9 (pmjectcd 7.0hnNE) (projected 8.0hnNE T4:

^-N sole fault 93

Woodside Shale through the Jurassic Twin Creek Limestone, and 4) from Jurassic Entrada Sandstone through Cretaceous sandstones, shales and conglomerates of the Indianola and Mesaverde Groups. Flat sections of the Charleston and Nebo thrusts are present in incompetent shales and evaporites of the Mississippian-Pennsylvanian Manning Canyon Shale- Mississippian Donut Formation (black shale), the Permian Park Qty Formation (interbedded dolomite and black phosphatic shale), Triassic Woodside Shale (thinly interbedded shale and sandstone), and the Jurassic Arapien Shale (gypsiferous shale and halite). The most prominent of these is the 35 km long thrust flat developed in Arapien evaporites (Fig. 3.2; Black, 1965; Smith and Bruhn, 1984; Lawton, 1985; Gallager, 1985). Thrusts of the C-N allochthon were subsequently folded and in some cases reactivated to accommodate slip at the base of the Santaquin culmination on the Uinta Basin-Moimtain Boundary thrust (Uinta BMB thrust) - the sole thrust of the Uinta allochthon (Figs. 3.2 and 3.3) (Constenius and Mueller, 1996). Tectonic transport of the Uinta allochthon, the lower horse of the Santaquin antiformal duplex, took place from Campanian through late Eocene time (ca. 82-40 Ma) based on sjmorogenic growth strata deformed by fold-thrust structures. Structures representative of this shortening episode are superposed on the C-N allochthon in the form of thrusts with west-vergence and fault-propagation folds which include the Diamond Fork anticline and Thistle Dome, and the Tie Fork syncline (Walton, 1944) a growth syncline paralleling the leading-edge of Uinta BMB thrust. The stratigraphic record of this deformation is preserved in progressive unconformities in strata ranging from Campanian Blackhawk 94

Fonnation and Castlegate Sandstone to Maastrichtian-early Eocene North Horn Formation, Eocene Flagstaff Limestone and middle Eocene Green River Formation. The most spectaciilar exposures of synorogenic growth strata are foimd adjacent to the Oil Hollow anticline, a large fault-propagation fold that developed in Campanian-middle Eocene time (Fig, 3.2), Near the Hunt Indianola #24-1 well (Fig, 3.1), the basal S5niorogenic cover (Blackhawk Formation) of the anticline is vertical or overturned and there is eastward thickening and progressive flattening of dip in successively younger units of the Castlegate Sandstone-Price River Formation-North Horn Formation which culminates in flat Ij^g lacustrine beds of the Hagstaff Limestone (Khin, 1956; Pinnel, 1972; Rimyon, 1977; Witkind and Weiss, 1992). A change in provenance of thrust-belt derived lithic clasts, from a spectrum of Proterozoic-Jurassic clasts to a single formation-source of Pennsylvanian- Permian quartzite clasts, is also an earmark of the "Uinta" phase of culmination growth, Castlegate-Price River and Currant Creek conglomerates are dominated by Oquirrh Group quartzite clasts derived from erosional stripping of the upper horse of the antiformal duplex. Dramatic east to west thinning of North Horn Formation (1500m to 240 m or less; Fig, 3,2) near Thistle, Utah is also evidence of progressive structural-stratigraphic growth related to ttirust-slip on the Thistle and Little Diamond Creek faults and the "Deer Creek thrust." Geometrically, the Charleston and Nebo thrusts display fault-bend fold geometries reflective of fault trajectories that rcunp through mechanically competent imits and are bed-parallel in intervening weaker shales or evaporites of the Mississippian Manning Canyon Shale, Permian Peirk Gty 95

Formation, Triassic Woodside Shale and Jurassic Arapien Shale. The most prominent tiirust flat in the Arapien Shale is about 35 km in east-west length and imderlies most of the C-N thrust salient (Figs. 3.2, 3.3, and 3.4). The large amount of horizontal translation across the Arapien bedding plane was accommodated by a complex series of triangle zones with passive roof thrusts (east-dipping backthrusts) (Jones, 1982; 19%) developed first in the Arapien Shale and later in the other major detachment units within the allochthon. The frontal Arapien triangle zone that formed in response to Nebo and Charleston thrust displacements was eventually cut and displaced by the Charleston thrust by Santonian-early Campanian time (ca. 84-82 Ma). Motions on the Charleston (late stage) and the blind Uinta BMB thrust systems did little to advance the thrust front of the C-N allochthon. The small amoimt of eastward hanging wall translation and resiiltant deformation is expressed by the ramp-on-ramp leading-edge Charleston thrust geometry and development of the Oil Hollow fault propagation fold. Significantly, because eastward translation of the thrust sheets weis stalled at the front, thrust-shortening was taken-up by internal contraction of the thrust sheet. This resulted in formation of east-dipping backthrusts in an intraplate triangle zone. Shortening was facilitated by the architecture of the Santaquin culmination that had imparted an east-dip to ttie main detachment units and folded thrusts. Passive roof Uirusts of the intraplate triangle zone and backthrusts utilized this hanging wall anisotropy and detachments formed in the Manning Canyon Shale, Park Qty Formation, Woodside Shale and Arapien Shale, and in places, the Nebo tiirust was reactivated. The map pattern of the Little Diamond Fork fault suggests that it vised bedding planes in the Park Qty Formation and Woodside Shale as slip surfaces, whereas surface and subsurface data indicate that the Thistle and Dairy Fork faults sole into the Arapien Shale (Figs. 3.2, 3.3 and 3.4) (Baker, 1972; Bryant, 1992; Witkind and Weiss, 1992). The "Deer Greek thrust" is considered to be a thrust that reactivated folded, east-dipping segments of the Nebo thrust and the through-going slip surface also detached bedding planes in the Marming Canyon Shale. It is intriguing that east-dipping segments of the Nebo thrust, reactivated as backthrusts, locally produced yoimger over older thrusting (Figs. 3.2 and 3.3). The most prominent example of stratigraphic attenuation related to this tmusual thrust style was mapped by Baker (1959; 1964) as the Deer Creek thrust where upper-middle Oquirrh Group rocks rest discordantly on Mississippian Manning Canyon Shale. Figure 3,2 shows Oquirrh Group in the hanging wall of the reactivated segment of the Nebo thrust resting on Mississippian strata (Great Blue Limestone), with complete attenuation of the Marming Canyon Shale (ca. 5(X) m thick). Other examples of yoimger-over older Nebo thrusting ("Payson and White Lake Hills" thrusts) were discussed by Brady (1965) and involve stratigraphic relations as extreme as Oquirrh Group on Cambrian. Similarly, exposures of Jurassic strata resting discordantly on early Permian rocks in the Red Rock Meadows area suggest that the passive roof thrust of the frontal triangle zone has evolved into a yoimger over older thrust (E. J. Steme, written commim., 1997; G.H. Wahlman, written commim., 1996). This unusual style of thrusting is a direct coi\sequence of the intraplate triangle zone that developed during late stage construction of the Santaquin 97 culmination where segments of the Nebo thrust were reactivated as part of a passive roof thrust above an intercutaneous wedge (Qiarleston and Uinta BMB thrust sheets) of a triangle zone (Jones, 19%). The large magnitude of eastward translation of the wedge (ca. 50 km) resiilts in stratigraphic attenuation across segments of the reactivated Nebo thrust because in these places the sense of hanging wall-footwall thrust displacement has been reversed (E.J, Sterne, written conunim., 1997), Compare the position of Pennsylvanian-Permian reference point "r" with respect to Proterozoic-lower Paleozoic rocks of the Nebo thrust hanging waU. Reversal of movement on the Nebo thrust segment shown in Figure 3.2 has evolved to a state where about 1300 m of Mississippian-Pennsylvanian strata are missing because the hanging wall flat in the Mississippian Maiming Canyon and part of the ramp in the lower Peimsylvanian Oquirrh Group have been consimied by thrust displacements. However, how can we determine if these truly were backthrusts, some with strata! attenuation, or if they were later reactivated as normal faults? The answer is two-fold. First, synorogenic strata of the Maastrichtian-early Eocene North Horn Formation and yotmger synorogenic units thin dramatically in proximity to Thistle and Little EHamond Fork faults indicating ttiat ttiey were contemporaneous with trusting. Second, there are neighboring imextended backthrust-structures such as Thistle Dome, Diamond Fork anticline and the thrust relations at Red Rock Meadows that capture the late-stage thrust style of tfie C-N allcchthon (Neighbor, 1959; Walton, 1959; Constenius and Mueller, 1996; Haldar, 1997), 98

3.4 RECORD OF CHARLESTON-NEBO COLLAPSE 3.4.1 Superstructure Collapse and Extensional Triangle Zone Extensional structures superposed on the C-N allochthon record a phase of late Eocene-middle Miocene tectonism, in which the sole thrust of the Santaquin cidmination was extensionally reactivated and motions on the thrust-wedge that had driven intraplate contraction were reversed. Westward translation of the thrust-wedge caused the superstructure of the culmination to collapse and passive roof thrusts and other thrusts of the pre­ existing intraplate C-N triangle-zone were transformed into normal faults. Significantly, the leading-edge of the C-N allochthon below the Arapien detachment surface (Dairy Fork fault) remained una^ected by normal-fault movements and the original thrust-fold architecture of the allochthon is preserved (Fig. 3.2). Two principle modes of extension are recognized in the C-N allochthon: 1) a network of half-grabens with up to 2 km of basin-fill are foimd superposed on the southern half of the allochthon; 2) in the northern part of the allochthon, a half-graben parallels the surface trace of the C-N sole thrust and the basin-bounding listric normal fault merges into the sole thrust. The southern half-grabens have bowl-shaped map patterns and profiles, exhibit stratal growth geometries in the synextensional basin-fill on both sides of the basin, and the basin-boimding faults are developed in imits that are considered to be mechanically weak. The bowl-shape geometry of the half-grabens was imparted by wedges of early Tertiary North Horn and Flagstaff strata in the hanging wall block. Normal faults that bound the half- grabens follow tiie same stratigraphic imits over long distances or cut discordantly across bedding to link segments of the bedding-plane faults. The 99

hanging wall geometry of the allochthon created an inclined stratigraphic column in which the Manning Canyon Shale, Park Gty Formation and tiie Arapien Shale, imits which had accommodated horizontal translation of the C-N allochthon over large distances, failed under extension. Unusual bedding geometries in the grabens, which heretofore were explained by models involving diapirism and depositional onlap (e.g., Witkind, 1987), are interpreted here as rollover anticlines and synextensional stratal growth formed by bed rotation above listric normal faults.

3.4^ New and Previous Concepts Half-grabens superposed on the southern part of the C-N allochthon encompass a region of complex geologic relations resulting from the overlapping effects of Cretaceous-middle Eocene fold-thrust shortening, late Eocene-middle Miocene extension, and Holocene dissolution of evaporites (Lawton, 1985; Constenius, 1995). Ardent proponents of a salt diapirism model (Witkind, 1983,1987, and 1992; Witkind and Page, 1983; and Runyon, 1977) considered many of the normal faults and folds to be the product of Jiurassic-Recent flow of Jurassic Arapien evaporite and shale. Geologic maps £md cross sections of the C-N allochthon by these authors show many of the half-grabens to be on trend with the axis of salt diapirs and that the allochthon rides above a salt mass several kilometers thick (Witkind and Weiss, 1991). Witkind (1987) also conceptualized an "erosional escarpment" model that sought to explain contact relations between the Maastrichtian- early Eocene North Horn Formation and older Paleozoic and Mesozoic rocks. The trend of this "escarpment" coincides witii tiie western margin of several 100

of the half-grabens. North Horn strata, as noted in Witkind's report, invariably dip away from the escarpment and are folded into a syncline that parallels the escarpment. Notice in Figure 3.5 that both grabens have similar hanging wall geometries even though ttie substrate beneath the North Horn unconformity is different (Jurassic Arapien Shale versus Triassic Woodside- Ankareh). This observation argues for a fold mechanism other than diapirism, because the Arapien Shale floors only one of the grabens. The unusual bedding geometries which have been explained by diapirism and escarpment models are viewed here as non-compressional folds formed by bed rotation above listric normal faults. The "erosional escarpment" is reinterpreted as a network of normal faults, based on recogiution of the following critical relationships: 1) listric normal faults bound tiie Little Qear Creek, Little Diamond Greek, Payson Lakes and Santaquin Meadows half-grabens, 2) exposures of the basin-boimding faults show them to be bedding-parallel faults with brittle fault-rock fabrics (gouge and breccia), 3) late Eocene-middle Miocene basin-fill strata of the Moroni Formation exhibit stratal growth; that is gradual flattening of dip in successively yoimger imits (McMechan and Price, 1980; Constenius, 1982), and 4) reverse-drag and rollover anticlinal structures (Xiao and Suppe, 1992; Schlische, 1995) are present in the hanging walls of the normal faults. Outcrop and substirface study of the North Horn Formation adjacent to the "escarpment" (i.e.. Little Diamond Fork normal fault) show that it was deposited on an unconformity surface with large paleotopographic relief. However, fault-proximal tilting and folding of these beds resulted from a deformational mechanism related to slip on normal faults and not to late 101

Figure 3,5. Geologic cross section C-C of Little Diamond Fork and Little Qear Creek half-grabens 5 km south of Thistle, Utah, shows bedding-plane normal faults in which the thrust-fold structural fabric of the allochthon has controlled the geometry of these basin bounding faults. The Little Diamond Fork fault and Thistle fault are extensionally reactivated bedding-plane thrusts in basal Triassic (Woodside Shale) and Jurassic Arapien Shale. The geometries of these faults have created unusual bowl-shape half-grabens with extensional-stratal growtii developed on both sides of the basins. Dating of tephras at base of Moroni Formation and a vertebrate fossil from upper Moroni (stratigraphic position indicated, not locality), suggest that collapse of Charleston-Nebo allochthon began in latest Eocene and continued until early or middle Miocene time. Sources of data: Harris (1953) and unpublished mapping by Consteruus (1994). 40Ar/39Ar age-data furnished by TP. Flood and P.W. Layer (written commun., 1997). See Figure 3.2 for abbreviations. Bowl-shaped half-grabens c c West Little Diamond Creek Little Clear Creek East half-graben half-graben Thistle Lower Moroni Fw Little Diamond fault U^erMoroniFm Fork fault 40.9+/-1.9 Ma fSOOO

-2000 Buried Hogback

-WOO

- s.l.

- -1000

-2000

oN> 103

episodes of thrusting, as proposed by Witkind (1987). The largest relief on the unconformity surface is foimd not where mapped by Witkind (1987) but to the east where ~700-800 m of relief are preserved along an ancestral Oetaceous Indianola Group hogback (Fig. 3.5). The Indianola hogback is partially exposed at the surface and is well delineated by reflection seismic data that show onlap and tnmcation of gently dipping to flat-lying lower North Horn strata (Haldar, 1997, Constenius, impublished data). Previous discussions have emphasized the role of the Arapien Shale as a detachment horizon that accommodated both shorting and stretching of the allochthon. As shown in Figures 3.2 and 3.5, the Arapien Shale swells from about 200 m to a thickness of -1.2 km, suggesting the possibility of salt- diapirism. However, the anomalous thickness of the Arapien Shale can be constrained to be pre-Maastrichtian because the basal North Horn unconformity bevels this unit. Borehole and seismic data from the Absaroka thrust sheet in norttieast Utah and southeast Wyoming have been used to show that the Preuss Formation (Arapien Shale stratigraphic equivalent) flowed during fold-thrust shortening. Preuss evaporites are commonly thin on the crest of anticlines and very thick on the forelimbs and backlimbs of anticlines and in synclines (Lamerson, 1982; Royse, 1993; Coogan, 1992; Constenius, 1996). The thickened Arapien Shale shown in Figure 3.5 resides on the forelimb of a large fault-bend fold. Thus, timing constraints and structural analogies indicate that the overthickened Arapien Shale is the product of salt-flow concurrent with shortening. The salt-diapirism and escarpment hypotheses are unsatisfactory explanations for the genesis of the grabens and masked the underlying cause 104 of the deformation - extensional collapse and westward transport of the thnist sheets. Salt related phenomena that post-date shortening and extension are restricted to an isolated ocoirrence of salt dissolution that has produced a large sinkhole (~120 x 300 m) about 4.6 km south-southwest of Thistle, Utah, dubbed Witkind's Revenge sinkhole (Constenius, unpublished mapping, 1994). Salt-diapiric folds imder the basins and great thicknesses of salt in the footwall of the C-N allochthon are not imaged on reflection seismic data (Fig. 3.4).

3.4.3 Structural Framework of Bowl-Shaped Half-Grabens Outcrop expression of the stratigraphic-structural control on the position of the basin-bounding normal faults is seen on maps of the area by Baker (1972), Pinnel (1972), Yoimg (1976), Runyon (1977) and Witkind and Weiss (1991). These maps show that the normal faults which boimd tiie half- grabens follow the same stratigraphic units over long distances (10-40 km) or cut at right angles to the stratigraphy to link segments of the bedding-plane faults. These units are the Manning Canyon Shale, Park Qty Formation- Woodside Shale and the Arapien Shale - the same mechanically weak beds which accommodated horizontal translation of the Qiarleston and Nebo thrusts over large distances. Locally, the G5^sum Springs Member of the Jurassic Twin Creek Limestone served as a detachment horizon for the Thistle fault 5 km south of Thistie, Utah (Fig. 3.5). Subsurface data from this area and geologic relationships in tiie northern part of the allochthon indicate that these normal faults sole into the C-N sole thrust which has been 105

extensionally reactivated (Fig. 3.4). The half-grabens tiiat formed in response to slip on the Little Qear Greek, Little Diamond Creek and other basin-boimding normal faults have "bowl-shaped* map patterns and cross sections, and exhibit strata growth geometries in the sjmextensional basin-fill on both sides of the basin (Figs. 33 and 3.6). For instance, Moroni Formation strata on the east side of the Little Qear Greek half-graben dip 45° W at the base and progressively flatten upsection to horizontal at the synclinal axis of the basin. Similarly, on the west side of the basin tiie same rock tmits dip eastward and flatten upsection. In contrast, half-grabens in the Q)rdilleran fold-thrust belt commonly are asymmetrical, with sjmclinal axes of half-grabens closely paralleling basin- boimding faults, and tmidirectional stratal growth geometries (e.g.. Smith and Bruhn, 1984; Janecke, 1992; Constenius, 1996, Mohapatra and Johnson, 1998). The "bowl-shaped* half grabens also differ from conventional half-grabens in that the master basin-bounding fault is often not in contact witii growth strata in the graben (e.g.. Fig. 3.5, Little Qear Greek half-graben). This unusual style of normal faulting and sedimentation is coi^dered to be the product of excess rock volume in the hanging wall of the faidt; that is deformed "wedges" of North Horn and Flagstaff strata in the hanging wall block. A discriminating characteristic of the Littie Qear Geek half-graben is its "U-shape" map pattern, a reflection of normal fault reactivation of the Arapien Shale detachment and an inherited synclirud fold geometry (Figs. 3.1 and 3.2). The Thistie and Dairy Fork faults, which originated as out-of- sjmcline tiurusts, sole into Arapien Shale evaporite units that have been folded into a north-plimging syncline. Back-sliding and resultant evacuation 106

Figure 3.6. Oblique areal photograph of bowl-shaped Little Diamond Fork half-graben 7 km north of Thistle, Utah, illustrating synclinal folding of pre- and synextensional strata. Conglomerate units in basal Moroni and Uinta formations stand as U-shaped ridges characteristic of the bowl-shaped half- grabens superposed on the Charleston-Nebo allochthon. See Figures 3.2 and 3.7 for abbreviations. View to west. Photo by Constenixis, October, 1996. Spanish Forit Peak 108

of the Santaqtiin ctdmination from this area caused its superstructure and the overlying sjoicline to collapse. Normal-slip on the Thistle and Dairy Fork faults, along with the Little Diamond Creek faiilt facilitated this extension and these motions were transferred down to the reactivated sole fault of the culmination. Notice in Figure 3.2 that the Thistle and Dairy Fork faults merge at depth and tfiat the Little Qear Creek half-graben appears twice because of the "U-shaped" plan of the half-graben. The timing and nature of normal faulting can be inferred from smrface relationships because S5mextensional stratal growth is present in volcaniclastic rocks of ttie late Eocene-middle Miocene Moroni Formation, and North Hom-Flagstaff pre-extei\sion strata have been folded into rollover style structures- These structural patterns represent the two modes of reverse drag that are symptomatic of listric normal faulting (Constenius, 1982). This indicates that displacement and rotation along the Little Qear Creek and Little Diamond Fork faults were synchronous with Moroni sedimentation. Potassium-argon dating of volcanic clasts and tephras from the Moroni Formation gives ages which mostly range from 38 to 33 Ma, but also include a single age of about 24.5 Ma (Witkind and Marvin, 1989). Oasts of pimiice from tiie base of the Moroni 5rield ^Ar/^^Ar integrated ages of about 38-40 Ma (Fig. 3.5). The age of the uppermost part of the Moroni Formation is about 16-20 Ma based on the recent discovery of Blickomylus sp., the first vertebrate fossil from this unit (M.R. Dawson, 1994, written communication). Age-bracketing of extensional reactivation of the allochthon based on these data is late Eocene-middle Miocene (ca. 40-16 Ma). This contrasts markedly with previous studies that considered normal faulting in this area to be 109

Neogene (ca. <25 Ma) in age (Royse, 1983; Hopkins and Bnihn, 1983; Houghton, 1986).

3.5 DEER CREEK DETACHMENT FAULT 3.5.1 Geologic Setting The sole fault of the C-N allochthon is exposed along the northern margin of the thrust salient and in places it has been extensionally reactivated by faults of the Deer Creek detachment fault system. Extensional reactivation of the C-N thrust along the Wasatch front was first identified by Christiansen (1950) who remarked 'Evidence is accumulating which indicates that the scarp of the central Wasatch Mountains between Draper and Alpine and

probably also north and south of this area is the stripped and slightly

modified sole of a major thrust zone. Stripping and laying bare of the thrust

surface was accomplished by late Tertiary erosion and probably by local

normal faulting on the thrust surface in a reversed direction from the thrust movement." Exposures of the Deer Creek detachment fault show it to dip 25- 50° south and map relations indicate that it merges to the east with the low- angle "thrust" faults or the main strand of the C-N thrust (Baker and Crittenden, 1961; Riess, 1985; Houghton, 1986). Two deformations conspired to fold and elevate the C-N sole fault ftrom depths of 10 km or more to the surface. The first was Ccmipanian-middle Eocene contractile growth of the Cottonwood arch-Uinta Moimtains uplift - elements of the Uinta allochthon. This was followed by late Eocene-Recent, southward and eastward tilting of the rock column related to development of the Cottonwood metamorphic core complex. Exhimiation of the C-N sole fault gives us a fortuitous down- 110

structure view (Mackin, 1950) of the thrust and later normal faults that reactivate the thrust surface. Down-structure views of the sole fault reveals that the Deer Creek detachment fault is superposed on a iriajor footwall ramp of the C-N thrust (Fig. 3.7) (Baker and Crittenden, 1961; Baker, 1964; Baker, 1972), The eastern limit of the detachment fault system is mapped near the west-end of the Arapien Shale thrust flat whereas the main site of extensional reactivation is found west of the thrust flat where the thrust ramps from the Mississippian Donut Shale through to the Jurassic Twin Creek Formation. Structures of the Cottonwood metamorphic core complex reside in an area formerly described as the Cottonwood arch; but the arch represents only one phase in the structural development of the central Wasatch Moimtains (Eardley 1939; Bradley and Bruhn, 1988) (Fig. 3.1 & 7). Here, a 6-8 km thick succession of Precambrian to Mesozoic rocks have been arched into a broad east-plimging anticline that has been tnmcated to the west by the Wasatch normal fault. At the core of the structure are Proterozoic crystalline rocks of the Little Willow Formation. Bisecting tiie core complex is the Wasatch igneous belt which is comprised of nine late Eocene-Oligocene stocks aligned in an east-west orientation and superposed on the Cheyenne belt crustal suture. U/Pb ages of zircons and other geochronologic data indicate that the stocks decrease in age from east to west (Fig. 3.7) (Vogel et al., 1997 and references therein). Volcanic rocks of the late Eocene-Oligocene Keetley volcanic field bury the plimging nose of the Cottonwood eirch and represent the eastern imbreached limit of the Wasatch igneous belt. The Little Cottonwood stock is the largest of the stocks and composed of quartz monzonite emplaced at deep 111

Figure 3.7. Geologic index map of Cottonwood metamorphic core complex and related structural elements. Emplacement ages for stocks based on zircon U/Pb analyses and location of dated sample shown (buUseye, site numbers are keyed to Table 2 in appendix). Sample sites for 40Ar/39Ar, K/Ar, and fission track geochronology designated with open circles. Triangles denote sample sites and paleodepth determinations in kilometers Qohn, 1989) (Site abbreviations: AF, American Fork Canyon (white line); BX, Box Elder Peak; DC, Dry Creek Canyon; JL, Jacob's Ladder; SL, Silver Lake Cirque; TT, Tibbie tephra; and, TVC, Tibbie volcaniclastic unit. Map adapted from Bryant (1992). Rock unit abbreviations not shown in Figure 3.2; Mdo, Mississippian Donut Shale; PPu, Pennsylvanian-Permian imdifferentiated; JTru, Jurassic-Triassic undifferentiated; and, Tt, late Eocene-Oligocene Tibbie Formation. Salt Lake City aaaaaaa/ Ontario stock 'Keetley volcanic'' (36.0 +/-10) Aha stock field X 5.0 */-10) Clayton Peak stock Wasatch fault (Salt Lake segment) sAfSO/ ^ . ^2^

Comer Creek / / / ^fV ///////, tectonite #>>vX VnV-^'Vj^wXS. 6 n ,

^"Little Condhwo'od^Jiir'^^'x^^i Traverse >10.4

/ f / / / /r\f^ • \ > \ N N S thrust footwall ramp

Deer Creek detachment fault /0ftm Silver Lake fault Wasatch fault Box Elder Peak anticline / Charleston-Nebo thrust (Provo segment) Tibbie half-graben 113 cmstal levels (ca. 6-11 km) (Parry and Bruhn, 1986; John, 1989). Deep-seated fault-rocks of the Comer Creek tectonite (cataclasite, phyllorute, breccia) form a 100 m thick tectonic carapace on the southwest margin of the Little Cottonwood stock (Bullock, 1958; Parry and Bruhn, 1986, 1987; Houghton, 1986). Study of the WIB stocks lead John (1989) to conclude that the igneous belt and its country rock have been differentially uplifted, imparting a 20° east-tilt to rocks of the core complex and that exposures of the crust ranging from 11 to 1 km paleodepths are preserved (Fig. 3.7). Consequently, present- day exposxires in the central Wasatch Mountains represent an oblique crustal slice that extends from near the paleosurface to deptiis that were near the brittle-ductile transition. Critical inspection of the geologic record in the southern part of the Cottonwood core complex in the vicinity of American Fork Canyon and along ttie trace of the Deer Creek detachment, provides the basis for doamienting, (1) late Eocene-early Miocene, extensional collapse of the C-N allochthon, (2) exhumation of mid-crustal rocks in the core of the Cottonwood arch, and (3) formation of brittle-ductile fault-rocks in the Deer Creek fault zone. Artifacts of this deformation are foimd in the structural style of the Box Elder Peak anticline, the Tibbie , slip direction indicators on the Deer Creek fault system, and faults previously mapped as thrusts with yoimger over older map relations.

3.52 Structural Evidence of Extension The interplay between thrust-fold and later collapse structures of the C- N allochthon is most dramatically seen in the walls of American Fork 114

Canyon (Fig. 3.7). Exposed here is the Box Elder Peak anticline, an immense rollover anticline (Gibbs, 1984; Ellis and McQay, 1988), broken at its crest by normal faults that sole into the Deer Creek detachment (Figs. 3.8, 3.9, 3.10, and 3.11). Geologic cross section A-A' illustrates that the Box Elder Peak anticline is a compound structure comprised of two superposed anticlines. Construction of tiie anticline took place in four phases. First, during the upper Jurassic-late Cretaceous (middle Campanian) (ca. 120-84 Ma) lateral ramps developed in the hanging wall of C-N thrust in a zone of oblique and left-lateral strike-slip faulting that characterize the north margin of the allochthon (Paulsen, 1997). The sole thrust was the progenitor of the Deer Creek fault, whereas the overlying imbricate thrust would later be used by normal faults boimding the American Fork crestal collapse graben. Second, late Campanian-middle Eocene (ca. 82-40) growth of the Cottonwood arch (Uinta allochthon) combined with the earlier stacking of imbricate thrusts, imparted south-dip on the rock column. Third, late Eocene-early Miocene (ca. 40-20 Ma) bed rotation associated witii transtensional slip on the curved Deer Creek fault surface created the north-dipping limb of anticline and resultant rollover structiure. Strata of the Tibbie Formation can be used to define this limb rotation episode as an extensional event because if they are restored to horizontal, the north-limb of the anticline is flattened or inverted and hanging wall rocks of the Deer Creek fault form a south-dipping monocline. Fourth, late Eocene-Recent (ca. 40-0) flexural-isostatic exhimiation tilted the entire rock column south and east, steepening the south-limb and east-plimge of the Box Elder Peak anticline. Study of the Tibbie Formation is central to imraveling the history of 115

Figure 3.8. Geologic cross section A-A' of Box Elder Peak anticline, a compound structure comprised of two superposed anticlines and the product of five structural phases: 1) lateral ramps developed in hanging wall of Charleston-Nebo tiirust in Jurassic-late Cretaceous time, 2) folding of the C-N allochthon during late Cretaceous (Campanian)-middle Eocene construction of Cottonwood arch (Uinta alochthon), 3) late Eocene-early Miocene extensional reactivation of C-N thrust plane by faults of the Deer Creek detachment fault system, 4) late Oligocene intrusion of Little Cottonwood stock, and, 5) late Eocene-Recent flexural-isostatic exhumation related to Cottonwood core complex. Hanging wall extension extension is tiered with slip along curved, reactivated sole thrust responsible for structural development of Box Elder Peak rollover anticline, and American Fork crestal collapse graben related to subsidiary normal faults rooted to reactivated C-N thrust imbricate. Based on mapping by Baker and Crittenden (1961). Subdivisions of Proterozoic Big Cottonwood Group (pCs) and Mississippian undifferentiated (Mu) abbreviated: Ybc, Big Cottonwood Formation; Ymf, Mineral Fork Tillite; Ym, Mutual Formation; Ml, Gardison Limestone, Deseret Limestone, and Humbug Formation; Mbg, Great Blue Limestone. Box Elder Peak anticline ^ American Fork A' South crestal collapse graben North American Fork Elevation Forest Gate Little Cottonwood rSOOO fault stock Box Elder Peak - 4000 Silver Lake Ridge

"2000 N / S• N / \ TVrn* » S• N / N NNNNlI NNS

Stiver Lake mylonite / / , , ^ < Deer Creek detachment fault system - i ^reactivated Charleston-Nebo thrust Alta thrust --woo

ON 117

Figure 3.9. Geologic cross section B-B' of Tibbie half-graben and Box Elder Peak anticline. Stratal growth in Tibbie half-graben establishes that the Box Elder Peak anticline is an extensional structure - rollover anticline. Restoration of basal Tibbie strata to their original subhorizontal attitude unfolds the northern limb of tiie anticline. Based on mapping by Baker and Crittenden (1961), Crittenden (1965), Paulsen (1997), and unpublished mapping by Constenius (1994). B Box Elder Peak Q' anticline Southwest ^ , Northeast American Fork crestal collapse graben Deer Creek aetachment fault system r 5000 Mount Timpamgos Stiver Lake Ttbble detachment '4000 half-graben - 3000 i.Uvi'.kK -2000 fm - 1000

i^iP -1000 119

Figure 3.10. Oblique arecd photograph of Box Elder Peak anticline and American Fork crestal collapse graben, American Fork Canyon, Utah. Folded imbricate thrust of Charleston-Nebo thrust system carrying Proterozoic and Cambrian strata seen in upper right. Form lines at photo center highlight fold hinge of Box Elder Peak anticline down-dropped along faults of the American Fork graben. See Figures 3.2 and 3.7 for abbreviations. View to west. Photo by Constenius, July, 1995. 120 121

Figure 3.11. Oblique areal photograph of American Fork Canyon, Utah, highlighting American Fork crestal collapse graben and truncation of Proterozoic and Cambrian strata by an imbricate of the Charleston-Nebo thrust. See Figures 3.2 and 3.7 for abbreviations. View to north. Photo by Constenius, July, 1995. 122 123

the Cottonwood core complex because contained within this unit is a record of the timing of extension, the level of pre-late Eocene hanging wall erosion, and the history of unroofing of the footwall of the Deer Creek fault. The synextensional nature of the Tibbie Formation is expressed by the systematic flattening and thickening (inferred) of its beds toward the Deer Creek detachment fault and by fault-proximal fades. Basal rocks of the Tibbie basin- fill assemblage dip up to 57° northeast (Baker and Crittenden, 1961; Constenius, 1994, impublished data) and progressively flatten upsection and toward the Deer Creek fault where they dip about 30° northeast (Fig. 3.12). The Tibbie Formation is a proximal alluvial fan deposit composed of brick red and grey, angular, boulder-cobble conglomerate, conglomeratic sandstone, mudstone and limestone. Paleocurrent directions determined from pebble imbrication indicate that Tibbie clasts were derived firom a rising footwall terrain to the northeast (Fig. 3.12). The composition of lithic clasts at the base of the Tibbie Formation record stripping of the eroded remnant of the C-N allochthon from the Cottonwood arch. Qasts fotmd in these lowest units of the Tibbie were predominantly derived from Proterozoic, Cambrian and Mississippian formations (site 1, Fig. 3.12). This initial level of footwall erosion is consistent with Tibbie subcrop mapping that indicates that the basal Tibbie imconformity had beveled the C-N hanging wall down to the Mississippian Great Blue Limestone, Mississippian Maiming Canyon Shale, and the lower­ most Pennsylvanian units of the Oquirrh Group (Baker and Crittenden, 1961; Baker, 1964). Less than 400 m upsection, the clast composition changes markedly as an influx of Pennsylvanian Weber Sandstone clasts denotes 124

Figxire 3.12. Paleocurrent directions determined from pebble imbrication fabrics (rose diagram) and coimts of lithic clasts (pie diagram) from five sites in the Tibbie half-graben. Proximal fan deposits of Tibbie Formation were derived from rising Deer Creek footwall to northeast. Tibbie strata rest unconformably on a Qiarleston-Nebo hanging wall that had been erosionally stripped to the level of the Mississippian Great Blue Limestone and Manning Canyon Formation, and tiie lower Oquirrh Group prior to late Paleogene extension. Provenance data suggest that the C-N thrust sheet was quickly breached in tiie Deer Creek footwall terrain, but that the level of Oligocene denudation was no deeper than the Pennsylvanian Weber Formation. Significantly, sites 1-4 contain no volcanic or plutonic clasts, indicating that early phase of Deer Creek extension (> 36 Ma) predated Wasatch igneous belt magmatism. 40Ar/39Ar age-data furnished by TP. Flood and P.W. Layer (written commim., 1997). I i5f| tU Wasatch fault

North

pCs-CD 10 km

Charleston-Nebo thrust Ppu(Pw)

Deer Creek detachment pCs-CD Tibbie half-graben

Ar/Ar .5 +/- 0.2 Ma .5 +/-1.4 Ma a = 50 / Fw/PPoq Ar/Ar 36.4 +/- 0.2 Ma 25" 45° '

205 n= 113 n=6l8 Pw/PPoq Porcupine Gulch dikes 218 II = 337 Silver Lake detachment fault 249 n = 62 Deer Creek detachment fault 1 km N» 126 breaching of the C-N thrust sheet (sites 1 & 4, Fig. 3.12). The lowest part of the Tibbie Formation is composed solely of clasts of sedimentary provenance and appears to predate nearby WIB magmatism (Fig. 3.12, sites 1-4). Although parts of the Tibbie Formation are within a kilometer of the Little Cottonwood stock, clasts of the intrusive or its contact metamorphic aureole are absent (Baker and Crittenden, 1961). However, the inception of WIB volcanism may be signaled by a single 2 cm thick tephra and rare tephra intradasts discovered 400 m upsection from the base of the Tibbie Formation (near site 4, Figure 3.12). Biotite extracted from this tephra yielded a ^Ar/^^Ar step-heat plateau age of 36.4 +/- 0.2 Ma. In contrast, the upper Tibbie basin-fUl assemblage contains a plethora of igneous cobbles and boulders derived from tiie Keetiey volcanic field (ca. 36-32 Ma) (Fig. 3.12, site 5). Volcaniclastic rocks of imknown provenance overlie the Tibbie Formation with angular imconformity. These rocks yielded a whole-rock K/Ar age of 27.5 +/-1.4 Ma and a hornblende ^O^r/^^Ar step-heat plateau age of 28.5 +/- 0.2 Ma (isochron age 28.8 +/- 0.2 Ma), which limits the age of the lower Tibbie basin-fill assemblage. Tibbie strata beneath the unconformity dip 50-57° northeast and tlie dated volcaniclastic unit dips about 25° northeast. This indicates that Tibbie strata have sustained a considerable amoimt of late Eocene-Oligocene and post-Oligocene rotation related to normal-slip on the Deer Creek detachment, respectively. The average rate of bed rotation based on these bedding orientations and the dated units would be about 30/m.y. Kinematic data from the Deer Creek fault zone indicate that the hanging wall of the C-N thrust was transported at least 5-7 km west- southwest in late Eocene-middle Miocene time (ca. 40-16 Ma); later during Ill

Basin & Range extension (ca. 17-0 Ma), western segments of the fault system were reactivated by the Wasatch normal fault system. Tracing the Deer Creek normal fault west from ihe Tibbie half-graben and Box Elder Peak anticline shows it to be collinear with what Parry and Bruhn (1986,1987) considered a transverse segment of the Wasatch fault (Stilt Lake-Provo segment bovmdary). We consider the normal fault that runs along the south margin of the Little Cottonwood stock to be the Deer Creek detachment fault (Crittenden et al., 1973; John, 1989), a normal fault that predated Neogene movement on the Wasatch fault. Fault-rocks of the Comer Creek tectonite in the footwall of the Deer Creek fault (including the supposed Wasatch fault segment) grade uniformly from brittle to brittle-ductile and ductile domains from east to west along the fault (i.e., shallower to deeper; ca. 5 to 11 km paleodepths) and there is no abrupt change at the Wasatch fault line (Figure 3,7; Houghton, 1986; Yonkee and Bruhn, tmpublished report, 1991). Based on the similarity of kinematic data from the Deer Creek and "Wasatch" faults, we concur with Houghton (1986), who proposed that these faults formed during the same tectonic event. In addition, physiographic and gravity modeling suggest that Deer Creek strand of the Wasatch fault is not the most active part of the fault system. The present Utah Valley floor is not located along the Deer Creek detachment, but rather, on the south side of the Traverse range (Fig. 3.1, 3, 7). The thickness of Miocene-Recent basin-fill (~2-4 km) to the northwest, west and south of the Traverse Range (Zoback, 1983) suggests that the most active parts of the Wasatch fault system border the outer margin of ttie range. In this context the Traverse Range is a mid-Cenozoic fault-block (horse) trapped 128

between the Deer Creek detachment fault to the north and the main strand of the Wasatch fault to the south. The Traverse range is not an isolated entity in this respect, because a similar feature, known as the Salt Lake salient, is found 35 km north (Constenius, 1996; Bryant, 1992). Measurements of brittle shear zones mainly in the footwall of the Deer Creek detachment fault combined with published kinematic data were used to establish the transport direction of the Deer Creek hanging wall. The attitude of the detachment fault and fault planes and striae measured by us from brecdated fault-rocks in the Deer Creek fault zone are summarized in Figures 3.13 and 3.14. Kinematic results derived from these data combined with observations from previous studies indicate a west-southwest extension direction for the Deer Creek detachment and adjoining parts of the Wasatch normal fault. Slip direction determinations based on brittle fault-rocks indicate west-southwest oblique slip, on a 25° to 5(P south-dipping fault plane (Fig. 3.14, Houghton, 1986; Zoback, 1989, Yonkee and Bruhn, unpublished report, 1991). Notice that the calculated extension direction is uniform along the trace of the fault. Slip direction indicators on ductile fault- rocks (phyllonite) from the Deer Creek fault are horizontal and west-directed for lineations and S-C fabrics plunge 20-25® southwest and 40® southeast (Parry and Bruhn, 1986; Houghton, 1986). Ductile fabrics in these rocks have been overprinted by brittle deformation. The magnitude of west-southwest transport of the Deer Creek hanging wall is considered to increase from east to west along the trace of the fault and to be at least 5-7 km based on restored Permian Park Gty-Triassic Thaynes cutoffs in the Cottonwood arch versus tiiose in a splay of tiie Deer Creek 129

Figure 3.13. Summary of detailed structural data from Deer Creek detachment - Wasatch normal fault system. Trend and dip magnitude of fault plane shown with arrows and values. Fault planes and striae measured from Deer Creek-Wasatch fault-rocks plotted as equal area projections using FaultKin 3.8a program courtesy R. AUendinger, et al., Comell Univeristy, Ithica, New York. Normal sense of shear assumed for striae. Hinges of sheath folds (triangles), foliation (open circles), lineation (grey dots) and great circle best fit of fold hinge distribution (grey line, pole-black square) for Silver Lake mylonite plotted at upper right. Dry Creek (DQ data provided by A. Yonkee (written commim., 1996). Sources of fault-dip data: Baker and Crittenden (1961), Baker (1964) and Yonkee and Bruhn (impublished report, 1991). Dry Creek (DC) shear zones (48) Equal Area Wasatch Fault

littlf Cottonwood / stoc'^ ^ Silver Lake mylonite (SL)

i i -l^tachment fault i system

Dry Creek Canyon (DY) shear zones (86) Corner Creek Equal Area tectonite

Deer Creek Pass (DP) Mill Fork (MF) Mul Canyon Peak (MC) shear zones (55) shear zones (72) shear zones (29) Equal Area Equal Area Equal Area

• "-rjfsi. ? ...

o 131

Figure 3.14. Summary of kinematic data from the Deer Creek detachment - Wasatch normal fault system indicating WSW extension direction. Magnitude of WSW transport of Deer Greek hanging wall is considered to be at least 5-7 km based on cutoffs of Permian Park Qty and Triassic Thaynes Formations (Royse, 1983). Data are plotted as focal mechanisms using Bingham statistic average of shortening (P) and extension (T) axes calculated using FaultKin 3.8a program. Paleostress directions determined by Yonkee and Bruhn (impublished report, 1991) shown as a3 values. Fault-rock fabrics for phyllonite in Fort Canyon from Parry and Bruhn (1986). Dry Creek (DC) (T-axis 253'® -3°) Equal Area Wasalch Faull N 10 km

Utuc Cottonwood^ Phyllonite in Fort Canyon (FC) o3 = 249",0\ L-S(270°.(f) a3 = 250",5" 'jVa^tes Fms) S-C(12(f,4

:n: m -Hiu

•iTi;Ifjljii?: rthj;;. Utf:::::]!?:

N) 133 detachment (Fig. 3.7) (Royse, 1983). The horizontal offset of Proterozoic Big Cottonwood Formation preserved as a fault sliver between the C-N thrust and the Deer Creek fault compared with downdropped and rotated parts of the same formation suggest a minimimi of 8 km of hanging wall transport. The amount vertical separation across the Deer Creek detachment can be gauged from offset of the "Eocene unconformity" (i.e.. Rocky Mountain erosion surface). In the hanging wall of the fault the imconformity is at the base of the Tibbie Formation at a present-day elevation of ca. 1100 - 2400 m (downdropped and rotated). Paleobarometric data from the WIB were used by John (1989) to estimate the intrusive depths of the stocks and the position of the Eocene-Oligocene land surface (i.e.. Eocene unconformity). The Little Cottonwood stock, the stock closest to the Tibbie half-graben was emplaced at a depth of 5-6 km, and the top of the stock is now about 1-2 km above the unconformity (ca. 3300 m). These relations suggest a miiumum of about 7 km of vertical offset for tiie fault at this locality (Houghton, 1986). Significantly, this segment of the Deer Creek fault was active oidy in late Eocene-eeirly Miocene time. The large amount of offset on the fault is also manifested in the juxtaposition of deep-seated plutonic rocks and intensely contact metamorphosed Mississippian footwall rocks with unmetamorphosed Paleozoic and Tertiary hanging wall strata along the north limb of the Box Elder Peak anticline (Fig. 3.8). Plutonic cover rocks in the Deer Creek fault hanging wall that have been displaced >7 km laterally and vertically are foimd in the Traverse Range where Bullock (1958) mapped brecdated Oquirrh Group, and hydrothermally altered late Eocene volcanic rocks (ca. 38 Ma; 134

Crittenden et al., 1973) and boulder-cobble conglomerate (Fig. 3.7). The "Eocene imconformity* preserved here at the base of the volcanic rocks is in fault contact with fault-rocks of the Comer Creek tectonite that formed at depths of 9-11 km. Geologic maps show the C-N thrust zone to be an interleaved network of low-angle faults that show both older over younger and younger over older relationships. Previous studies (Baker, 1959; Riess, 1985, in part) attribute the origin of these thrusts to beheading of preexisting thrust-fold structures. Three factors preclude this assessment, however. First, synextensional strata of the Tibbie Formation are in contact with, or parallel, the trace of these low-angle faults (Fig. 3.7) (Baker and Crittenden, 1%1; Baker, 1964; Riess, 1985). Second, hanging wall Triassic-Permsylvanian rocks form a structural horse in this extensionally reactivated fault zone which has been displaced -7 km west from a footwall cutoff position (Royse, 1983). Third, "thrust" faults of the Alta thrust zone mapped in the Porcupine Gulch and Silver Lake areas, which we recognize as the Silver Lake detachment fault (new name), attenuate the stratigraphic section by 800-1000 m (lower Mississippian rocks missing) and cross-cut Oligocene latite dikes that were intruded during an extensional regime (Fig. 3.15) (Crittenden, 1965; Constenius, 1996; Vogel et al., 1997). The eastem part of the Silver Lake detachment is superposed on thrust-imbricated Cambrian-Mississippian strata of the Alta thrust system, and it is likely tiiat the detachment represents extei\sioi\ai reactivation of this earlier contractional fault system. The timing of detachment faulting is clear, because dikes with geochemical affinities to the Alta stock (ca. 33-35 Ma) are 135

cut by the detachment fault (Figs. 3.9 and 3.15) (Crittenden, 1965; Vogel et al., 1997). This contrasts with other parts of the Cottonwood core complex where numerous WIB dikes cross-cut the Alta thrust (Crittenden, 1965; Baker, 1966). Impressive cross-cutting relationships exposed in the west wall of Silver Lake cirque also define the Silver Lake detachment as an extensional structiure and give an encapsulated temporal view of the complex, rapidly evolving interactions of WIB magmatism and detachment faulting (Figs. 3.15 and 3.16). The Silver Lake detachment at this locality has structurally-stripped cover rocks of the Little Cottonwood stock and in places the upper margin of the stock itself. Fortuitously, preserved beneath ttie detachment fault is a smaU half-graben and an earlier formed mylonitic fault zone that records ttie earliest phases of extensional denudation of the Little Cottonwood stock. The Silver Lake mylonite (new name) is a 0.5 to 5 m thick south-dipping ductile shear zone that separates quartz monzonite of the Little Cottonwood stock from contact metamorphosed and tectonized limestone and dolomite (marble) of the Mississippian Donut Formation (Crittenden, 1965; Burge, 1959). The character of the shear zone changes markedly along its trace varying from light-green mylonite with sheath folds to sheared and altered igneous protolith interleaved with marble or quartzite. Stratal omission associated witii the Silver Lake mylonite is at least 2000 m based on the absence of rocks ranging from upper Proterozoic Big Cottonwood Formation to Mississippian Humbug Formation. The distribution of sheath fold hinges combined with the few lineation measurements suggest an east-west, or perhaps a more southerly extension direction (Fig. 3.13). The Silver Lake mylonite was later cut by a high-angle, northeast 136

Figure 3.15. Geologic map of Silver Lake detachment fault system showing location of dikes cross-cutting or cut by detachment fault. See Figure 3.7 for location map. Dikes shown as heavy black lines. Sample sites for 40Ar/39Ar, U/Pb, and fission track geochronology designated with open circles. Modified from Baker and Crittenden (1961), Crittenden (1965), and impublished mapping by Constenius (1996). 137

3199 m Uttte Cottonwood stock

Figure 16 Aua -r , ; Porcupine / / 36 thru^ j- 20^ >:Gulch Silver 3

Folded dike 292 m

ii-^SuasLQi^ 111 37^' 49 3(r-

Deer Creek normal fault Tibb^leTibBle half-grabenhalf-eraben *• / R ?•"" J' 1 138

Figure 3.16. Sketch of west wall of Silver Lake cirque illustrating cross cutting relations which evolved rapidly during coeval magmatism and extension about 30 Ma. The vestage of mylonite preserved in the small half-graben records an early phase of extension, the record of which was truncated elsewhere by displacements on the Silver Lake detachment fault. The beheaded half-graben and detachment fault were later intruded by three aplite dikes of the Little Cottonwood stock magmatic system. Sample sites and results from 40Ar/39Ar, U/Pb, and fission track geochronology shown. Apatite and zircon fission track data furnished by P.J.J. Kamp and P. A. Armstrong (written commim., 1997). Argon spectra plotted using MacArgon 5.11 progreim courtesy G.S. Lister and S.L. Baldwin, Monash Univeristy, Melbourne, Australia, and University Arizona, Tucson, Arizona. STRUCTURAL RELATIONSHIPS EXPOSED IN WEST WALL OF SILVER LAKE CIRQUE South North Silver Lake c , , , , Dikes detachment fault mylonite Uttle Cottonwood stock Marble tectonite U/Pb Zircon 30.5 +/- 0.5 Ma Fission Track Zircon 29.3 +/-1.2 Ma Fission Track Apatite 12.3 +/-1.5 Ma

-30 *»f| . ////// \ \ N \

K-Feldspar

•e 20- Aplite Dike _ Whole Rock Fraction ^®Ar Released rt 15 •' g; 0.0 0.2 0.4 0.6 0,8 1.0 20- Mylonite < Fraction Ar Released Whole Rock 15 -I r- -i 15 0.0 02 0.4 0.6 0.8 1.0 Fraction 39'Ar Released 140

trending normal fault, down-dropping the fault zone and the overl)nng marble tectonite by about 100 m. The resultant half-graben was beheaded shortly thereafter by the Silver Lake detachment fault. Although the Silver Lake mylonite and Silver Lake detachment appear to be different faults, it is likely that they are part of the same fault system and collectively form and extensional duplex. Early normal-slip on ttie detachment synchronous with main-phase intmsion of the Little Cottonwood stock formed a mylonitic floor fault etched on the stock. Downdropping of the mylonite caused the floor fault to be abandoned and a new fault path formed a roof thrust, the SUver Lake detachment. Aplite dikes emanating from the Little Cottonwood stock intruded or cross-cut these earlier formed structures and provide a relative age-limit to the Silver Lake detachment fault system. The results of U/Pb and 40Ar/39Ar experiments, which will be discussed more fully next, suggest that the minimimi age of slip on the Silver Lake detachment fault is about 27-28 Ma, and that the tectonic and magmatic relationships exposed in the cirque wall took place over less than 2-3 m.y. Lastly, the complex geologic relationships seen in the Silver Lake cirque are the product of two overlapping and competing extensional systems active during the formation of the Cottonwood metamorphic core complex. One oriented northwest-southeast controlled the orientation of WIB intrusive rocks (e.g., aplite dikes) (Olig, 1989; Vogel et al., 1997) whereas the other, the product of orogenic collapse, was directed west-southwest and responsible for creation of the Deer Creek and Silver Lake detachment faults. 141

3.5.3 Geochronologic Evidence of Footwall Exhumation Structural observations and dating of synextensional basin-fill chronicle late Eocene-middle Miocene motions on the Deer Creek detachment and resultant renewed hanging wall deformation of the C-N allochthon. At this same time, the footwall of the Deer Creek fault was tectonically denuded and thermally quenched simultaneous with intrusion of WIB plutonic rocks. Recognizing that thermochronologic study of the denuded Deer Creek footwall could yield insights not only related to the time of slip on the Deer Creek detachment system but also on the growth of the Cottonwood metamorphic core complex, we studied the exhumation history of the WIB with emphasis on the Little Cottonwood stock. Reconstruction of P-T-t (pressure-temperature-time) paths recorded in footwaU rocks and fault- rocks of the Deer Creek fault system, used: (1) the high closure temperature- based system of U/Pb zircon dating to determine the intrusion-age of the Little Cottonwood and other WIB stocks (Hodges, 1991), (2 ) synthesis of previous K/ Ar and fission-track studies, (3) results from new fission track dating of apatite and zircon (lower closure temperatures), (4) ^Ar/ 39Ar K- feldspar and whole rock step-heat experiments and diffusion modeling from the Little Cottonwood stock quartz monzonite and Comer Creek fault-rocks, and (5) paleodepth estimates from Parry and Bruhn (1986; 1987) and John (1989). Intrusive ages of four WIB stocks. Little Cottonwood, Alta, Qayton Peak and Ontario, each representative of different paleodepth levels, were estimated using U/Pb dating of zircon crystals (closure temperature > 700 o C; Hodges, 1991). Laboratory methods used in this analysis and a data table 142 summarizing analytical results are shown in Table 4.1. The intrusive ages for the four stocks determined from this experiment are shown in Figures 3.7 and 3.17. These results are consistent with previous studies by Crittenden et al., (1973) and Bromfield et al., (1977) and in the cases of the Little Cottonwood, Alta (Brighton) and Qayton Peak stocks, form a firm basis from which to compare other geochronologic data. Reiterated in this dataset is the younging of the WIB stocks from east to west (Fig. 3.7). The 30.5 +/- 0.5 Ma age of the Little Cottonwood stock is slightly younger but within analytical error of the previous K/Ar hornblende age of 31.1 +/- 0.9 Ma. Zircon mineral separates from the stock samples contained a large number of inherited grains and even though we were attempting to date pristine crystals produced in WIB magmas some inherited grains were analyzed. A positive side-effect of dating zircons with inheritance however, was the suggestion by the small number of discordant ages that the sources and migration paths of WIB magmas were superposed on the Wyoming-Yavapai crustal boimdary. We speculate that three discordant age groups and their protolith are identified, Archean Wyoming province (ca. 2518 Ma), Proterozoic Yavapai (ca. 1714), and Uinta aulocogen (ca. 1044-1111 Ma). Excellent access to the central part of the Cottonwood arch and a long history of mining have resulted in numerous geochronologic studies of the WIB (Crittenden et al., 1973; Bromfield et al., 1977; Morrissey, 1980; Naeser and Bryant, 1983, Bryant et al., 1989; Kowallis et al., 1990; John et al., 1997; Vogel et al., 1997). The results of these studies are summarized in Table 4.2, and two main conclusions can be drawn from this tabulation of WIB apparent ages. First, inspection of apparent ages from the Little Cottonwood 143

Figure 3,17. U-Pb concordia diagrams of zircon grains in the Little Cottonwood, Alta, Qayton Peak, and Ontario stocks. Analyses vary from single zircon grains to multiple grains (see Table 3.1). 144

LITTLE COTTON"WOOD STOCK: LITTLE COTTONWOOD STOCK 3;; Silver Lake Snowbird & Silver Lake Composite 3< CD mCM 30.5 + 0.6 Ma « CO o LITTLE COTTONWOOD STOCK CM Snowbird 31

30.5 ± 0.5 Ma 207V235 30.5 + 0.8 Ma

ALTA STOCK ALTA STOCK Brighton Guardsman Pass 40.

33.5 ± 1.0 Ma 35.0 ± 2.0 Ma (?) 207V235 207*/235

CLAYTON PEAK STOCK ONTARIO STOCK 3i

00 TO m CM « oCO CM 35.5 ± 1.5 Ma 36.0 + 2.0 Ma 207V235 207V235 145

Grain Apparent ages (Ma) characteristics

Size # Wt. Pbc Pb U 206m 206c 206 206' 207' 207' (u) (ug) (pg) (ppm) (ppm)

Little Cottonwood stocK - Snowbird, Utah (40° 35' 01" N, 111° 39' 23" W)

145-175 1 38 36 4.02 636 220 250 2.8 30.1 ±0.4 30.8 ±2.7 84±200

145-175 1 43 21 8.72 1748 1120 1468 6.7 30.7 ±0.2 30.8 ±0.5 43 ±37

145-175 1 33 22 4.51 849 395 499 4.8 30.1 ±0.4 29.8 ±1.5 2±110

145-175 1 28 30 6.52 1198 355 419 4.7 30.3 ±0.3 29.4 ±1.6 -40 ±130 Little Cottonwood stock - Silver Lake cirque A (40° 31' 24" N, 111 ° 41'01" W)

175-250 1 50 44 7.31 1357 485 545 4.9 30.6 ±0.3 30.3 ±0.9 9±60

145-175 1 19 10 5.64 911 382 570 2.4 30.5 ±0.6 30.6 ±1.1 44±73

145-175 4 76 20 3.62 692 820 1087 4.7 30.4 ±0.3 30.5 ±0.5 36 ±33 145-175 4 71 39 5.51 1051 595 679 5.1 30.2 ±0.2 30.4 ±1.0 43±71

Little Cottonwood stock - Silver Lake cirque B (40° 31' 24'' N, 111° 41' 01" W)

145-175 1 29 17 14.91 3009 1530 2137 6.6 30.6 ±0.2 30.6 ±0.5 33 ±30

175-250 1 65 64 5.67 963 312 337 3.6 30.5 ±0.3 30.3 ±1.9 18 ±140

145-175 4 91 41 2.65 449 320 360 3.3 30.2 ±0.3 29.8 ± 1.2 -7 ±92 145-175 4 89 17 1.13 210 350 480 4.4 30.3 ±0.5 30.4 ±1.2 34±81 Alta stock - Brighton, Utah (40° 35' 58'' N, 111° 35' 08* W)

145-175 1 42 73 105.89 489 22800 39409 9.4 1218±7 1407±10 1706 ±6 125-145 4 37 20 4.63 709 457 599 2.8 33.4 ±0.4 33.3 ±1.0 30 ±62

125-145 4 63 20 1.03 152 184 235 3.1 33.3 ± 0.9 33.3 ± 2.2 34±140

145-175 1 31 60 1.93 325 345 625 3.9 33.5 ± 0.9 33.5 ±1.3 33±60 125-145 4 38 64 5.04 912 1080 1974 5.5 33.5 ± 0.3 33.3 ± 0.5 20 ±26

Alta stock - Guardsman Pass(40°36'35" N, 1110 33' 38" W)

125-145 1 24 36 79.03 515 3230 3754 8.1 890 ±10 932±11 1033 ±9 80-125 4 45 61 3.49 367 113 120 1.8 34.8 ± 0.6 34.4 ±2.7 6 ±170

80-125 4 27 78 8.76 417 1150 1924 6.5 123±1 250 ±3 1721 ±10

80-125 4 31 21 1.41 139 95 113 1.6 35.5 ±1.9 35.6 ±3.6 47 ±190 80-125 4 32 25 10.26 254 720 1186 3.3 213±3 288 ±5 944 ±22

Clayton Peak stock - Guardsman Pass (40° 36" 2l" N, 111o 33" 18" W)

125-145 1 25 20 7.41 1050 478 631 2.5 35.0 ±0.5 35.4 ±1.0 59±58 146 1 M + 0 C 145-175 4 127 61 4.55 602 470 510 2.4 36.5 ± 0.5 36.2 ±1.1

125-145 4 68 21 5.72 824 950 1236 2.6 35.4 ±0.3 35.6 ±0.8 45±45

145-175 4 77 99 4.74 516 162 169 1.8 36.0 ±0.4 36.2 ±2.4 44 ±150

125-145 4 49 66 3.91 541 620 1686 2.1 35.4 ±0.7 35.4 ±0.9 1+ Ontario stock

125-145 1 18 16 12.60 445 785 1134 4.8 153±1 399 ± 5 2300 ±14

80-125 4 19 15 1.87 230 122 163 2.1 34.5 ±2.0 34.2 ±3.6 18 ±200 O O O + 1 C C C 80-125 4 22 13 2.23 166 212 323 2.8 85±5 644 ±94

40-80 4 10 13 3.99 538 175 265 2.5 35.4 ±1.7 35.3 ±2.6 27 ±130

40-80 4 9 19 35.30 3174 1010 1369 5.3 65±1 83±1 615 ±19

40-80 16 44 16 0.73 112 330 1696 3.1 36.6 ±1.9 36.6 ±2.2 35 ±70

80-125 8 34 24 7.80 361 650 1086 5.1 125±2 173±3 891 ±22

80-125 8 41 14 8.52 364 1380 4837 4.2 127±2 323 ±4 2186 ±5

80-125 8 35 16 46.68 483 8500 142670 5.4 519±2 1048 ±5 2414 ±5

Note: U-Pb samples were processed using a jaw crusher, roller crusher, Wilfley table, heavy liquids, and Frantz

magnetic separator. The non-magnetic zircons were sieved into size fractions and then selected (or analysis based

on optical properties using a binocular microscope. An effort was made to select grains with few fractures, inclusions and cores, and highly elongate rod-shaped crystals. The zircons were dissolved in HF>HN03 in 0.01 ml

Teflon microcapsules within a 125 ml dissolution chamber during a period of 30 hours at 245° C. The solutions were evaporated to dryness. 205pb/235-233u gpji^g ,^35 added, and the precipitate was dissolved in the dissolution

chamber in 3.1 N HCI for 8 hours at 225° C. Isotope analyses were conducted with a VG-354 mass spectrometer equipped with six Faraday collectors and an axial Daly detector. The measurements were made in computer-

controlled dynamic mode, with the Daly detector used simultaneously with the Faraday collectors to measure

^'^Pb. The gain factor of the Daly detector was determined continuously by comparing ^°®Pb(Faraday)/^°®Pb(Faraday) witti ^°®Pb(Faraday)/^°®Pb(Daly) and 2°®Pb(Faraday)/^°^P'5(Faraday) w'tti 206pb(Faraday)/^°^P''(Daly)- Isotopic data were processed utilizing data reduction and plotting programs of

Ludwig (1991,1992).

t Ptic - total common Pb in picograms. § Asterisk (*) denotes radiogenic Pb. # 206m/204 is measured ratio, uncorrected for blank, spike, fractionation, and initial Pb. ft 206C/204 and 206/208 are conrected for blank, spike, fractionation, and initial Pb. §§ Pb and U concentrau'ons have uncertainties of up to 25 percent due to uncertainty in grain weight. ## Constants used: X235 = 9.8485 x lO-""®, X238 = 1.55125 x lO'lO. 238/235 = 137.88. ttt All uncertainties are at the 95% confidence level. §§§ Pb blank generally ranged from 2 to 10 pg. U blank was less than 1 pg. ### Isotope ratios are corrected for fractionation of 0.14i 0.10%/amu for Pb and 0.20 - 0.40 for U02- tltt Initial Pb composition Interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206/204,0.3 for 207/207, and 2.0 for 206/208. §§§§ All analyses conducted using conventional isotope dilution and thermal ionization mass spectrometry, as described by Gehrels and others, 1991. 147

TABLE 4.2. SUMMARY OF WASATCH INTRUSIVE BELT AGE DATA

Lim J. COTTONWOOD STOCK - NW t rm F. COTTONWOOD STOCK - NF. FT APATITE 85 *!- 1.0 rai FT APATTTE 10.2 »/- 1.2 lAKC FT ZIRCON 245 0.6 #CR FT ZIRCON 29.7 •/- 1.9 lAKC FT SPHENE 24.1 •/- 0.6 #CR K/Ar BIOTTTE 273 •/- 0.9 ICR K/Ar BIOTITE 25.1 0.8 #CR K/Ar BIOTTTE 27.99 -/- 036 IJ K/Ar BIOTITE 24.61 •/- 054 #J K/Ar BIOTTTE 29.10 •/- 3.40 IJ K/ArBIOrrTE 25.15 •/-0.24 « *°Ai/^'Ar BIOTTTE 31.8+/-23 PLATEAU IF K/Ar BIOTITE 25.83 •/- 0.70 « K/Ar HORNBLENDE 3153 •/-4.40 IJ •^OArf^'Ar BIOTITE 28.4 •/- 05 PLATEAU #F U/Pb ZIRCON (4) 305 •/- 0.6 IC •"•Aif^'Ar HORNBLENDE 303-<-/-2.2 PLATEAU #F K/Ar HORNBLENDE 31.1 •/-0.9 #CR LITTLE CQTTQNWOQP STOCK • ^VHITE ECiE K/Ar HORNBLENDE 29.93 -•/- 1.66 #J K/Ar MUSCOVTTE 255 •/- 0.8 ICR K/Ar HORNBLENDE 30.72 •/- 058 #J K/Ar SERICTTE 23.41 •/- 0.46 IJ RWSr 24 *!- 6 #C1 t lTTIJ-: COTTON^VOOD STOCK - SF. LITTT J. COTTONWOOD STOCK - S^' FT APATTTE 4.9 •/- 0.8. 7.9 »/- 1.2 & 123 »/- 15 lAKC FT APATITE 6.4 •/- 2.0 lo 10.8 *!• 2.8 #K FT ZIRCON 29.8 •/- 1.2 lAKC FT ZIRCON 9J •/- 0.8 to 20.4 •/- 1.6 /K ^''Ai/^'Ar WHOLE RX 27.0+/-03 PLATEAU [DIKE] IF K/Ar SERICITE 17.6 •/- 0.7 t9 WHOLE RX 27.1 W-O.l PLATEAU[MYLONTTEJIF *°Ai/^'Ar K-FELDSPAR 7-17 SPECTRA (li8 •/-0.6 INT) #C K-FELDSPAR 12-27 SPECTRA (2i8 •/-13 INT) 0C '*°/W^'Ar K-FELDSPAR 12-40 SPECTRA #C *°Ai^'/Vr K-FELDSPAR 20-40 SPECTRA IC "^OAI/^'AT BIOTITE 243 •/- IJ PLATEAU #F •^^AJ^'AT K-FELDSPAR 28.6 •/-0.4 INTEGRATED IC ^°A«/"Ar BIOTTTE 283 •/- 15 PLATEAU IF ALTA STOCK . F^inORANULAR U/Pb ZIRCON (8) 305 •/- 05 IC FT APATITE 215 *!• 2.0 #AKC FT APATTtE 34.3 •/- 15 #M ALTA STOCK - roRPHYRY FT APATITE 35.6 •/- 3.6 #€4 '*°Arf^'Ar K-FELDSPAR 34.4+/-03 INTEGRATED IC FT ZIRCON 35.0*/-2.9 #AKC K/Ar BIOTTTE 31.7 *1- 1.0 IB FT ZIRCON 33.7 •/- 0.9 #CR K/Ar BIOTTTE 31.6*/- 1.0 IB FT SPHENE 32.7 •/- 0.9 #CR K/Ar BIOTTTE 3il •/- 1.2 IB K/Ar BIOTITE 32.1 •/- l.O #CR K/Ar BIOTTTE 314 *1- 1.0 ICR K/Ar BIOTTTE 32.6 *1- 1.0 fCR K/Ar BIOTTTE 31.17 V. 054 IJ K/Ar BIOTITE 30.17 •/-054 tS K/Ar BIOTTTE 3X8 •/- 0.7 ICR "•"ai^'AT BIOTITE 31.97*/-0.23 INT #J K/Ar BIOTTTE 33.26 •/- 055 IJ "'^At^^'Ar HORNBLENDE 33.72 +/- 050 INT #J •*°Ai/^'Ar BIOTTTE 32.28 ••/-0.24 INT IJ K/Ar HORNBLENDE 335 •/- 1.0 #CR K/Ar HORNBLENDE 35.1 *1- I.l ICR K/Ar HORNBLENDE 30.9 •/-15 ICR ^°Ai/^'Ar HORNBLENDE 29.90 •/-0.41 INT IJ U/Pb ZIRCON (4) 335 */- 1.0 #C CLAYTONFEAKSTOCK ALTA STOCK - GUAKDSMAN FT APATITE 16.9 +/- 2.0 #AKC FT APATTTE 175 •/- 1.9 lAKC FT ZIRCON 35.9 •/-3.1 #/VKC FT ZIRCON 37.2 •/-3.0 lAKC FT APATITE 41.2 •/-3.0 ICR ••"/W^'Ar K-FELDSPAR 34.4 •/-03 LNTEGRATED IC FT ZIRCON 40.9 •/- 2.2 »C8 K/Ar BIOTTTE 31.7 •/- 1.0 IB K-FELDSPAR 31-40 SPECTRA #C K/Ar BIOTTTE 31.6+/- 1.0 IB K/Ar BIOTITE 32.9 •/- 1.2 #B K/Ar BIOTTTE 3il +/- 1.2 IB K/Ar BIOTITE 34.8 •/- 1J #B K/Ar BIOTTTE 31.17+/-054 IJ K/Ar BIOTTTE 36.7 */• 15 #CR U/Pb ZIRCON (2) 35+/-2.0 IC ^°Ai:r''Ar BIOTITE 32.03 */-0.21 INT « BIOTTTE 3130-/-2.70 PLATEAU #F U/Pb ZIRCON (5) 355 •/- 15 »C PINR CREEK STOCK FT APATTTE 283 */- 6.6 #AKC FT /^ATTTE 38.9 +/- 2.9 lAKC FT APATITE 35.1 •/- 3.0 #M FT ZIRCON 35.7 +/- 2.0 lAKC K/Ar BIOTTTE 35.2 •/- 13 #B K/Ar PLAGIOCLASE 35.6+/- 13 (DIKE) IB K/Ar BIOTTTE 36.8*/-1.1 ICR K/Ar PLAGIOOASE 35.7 *1- 13 (DIKE) IB K/Ar BIOTTTE 35.63 *1- 0.32 « K/Ar BIOTTTE 33.4 +/- 13 IB K/Ar HORNBLENDE 41.08 */- 052 « K/Ar BIOTTTE 34.0 +/- 1.1 IB "•"Arf^'Ar HORNBLENDE 385 V-0.7 PLATEAU; K/Ar BIOTTTE 33.7 +/- 13 IB ISOCHRON 38.0 •/- 0.9 #F K/Ar BIOTTTE 34.5 +/- 1.4 IB K/Ar BIOTTTE 333 +/- 13 IB ^''Aj/^'Ar HORNBLENDE 36.6 +/- Z1 PIATEAU IF VALEO STOCK U/Pb ZIRCON (8) 36.0+/-2.0 IC K/Ar BIOTTTE 34.6 •/- 1.6 #B K/Ar HORNBLENDE 403 *1- 1.6 #B6 FLAGSTAFF STOCK K/Ar HORNBLENDE 39.8 •/- 1.2 #B7 K/Ar HORNBLENDE 39.7 +/- 1.2 IB K/Ar HORNBLENDE 37.8 +/- 15 (Dike) IB MAYFI^WKR STOCK HORNBLENDE 40.8 +/-1.9 PLATEAU IF FT APATTTE 38.9 •/-55 #M FT APATTTE 38.6 •/- 3.6 #/UCC FT ZIRCON 3Z5 •/-il #AKC PARK PREMIER STOCK K/Ar HORNBLENDE 41.2 •/- 1.6 #B Klfa BIomTE 33.9 +/- 1.2 IB ^°Ai/^'/Vr HORNBLENDE 32.7 •/- 1.4 PIATEAU IF K/Ar BIOTTTE 31.21 +/-0.80 IJ K/Ar HORNBLENDE 35.2 +/- 1.0 IB K/Ar HORNBLENDE 3Z26 +/- 0.27 IJ 148

INDIAN HOI.IjnW Pt.lTf; K/Ar HORNBLENDE 36J »/- 1.3 #B CI JXCOE STOCK K/Ar HORNBLENDE 36.1 •/- 1.3 #B FT APATTTE 30.3 »/- 1.0 #M KKF.TT.KY VOI/-ANlr.S TRAVERSE VQLCAiMCS K/ArBIOnTE 35.1 •/- 1.1 #CR K/Ar BIOTTTE 38J •/- l.l ICR K/ArBIOTlTE 32.7 •/- 1.0 #CR K/ArBIOnTE 34.0 *!- 1.0 ICR TIBBLE FORMATION K/Ar BIOTTTE 33.9 •/- IJ #B •^"Arf^^Ar BIOTTTE 36.4 •/-0.2 PLATEAU (LOWER TIBBLE] #F K/Ar HORNBLENDE 36.4 •/- 1J #B "•"Arf^'Ar BIOTTTE 28.5 •/-0.2 PLATEAU [LTPERTIBBLEJ #F ^"Arf^'Ar HORNBLENDE 38.S»/-11 PLATEAU #F VERTEBRATE FOSSILS -34-37 #N

IAMPROITF Sn J .S fMOON CYN • FRANnS. im WHITES CREEK DIKE (E. CRANDAU. Cm ""at^'AT PHLOGOPTTE 35.21 •/-0.10 PUVTEAUfMB "•"ai/^'AT PHLOGOPTTE 13.93 *!- 0.09 PLATEAU #MB "•"aj^'AT PHLOGOPTTE 34.93 •/-0.12 PLATEAU #MB "^^Ai/^'at PHLOGOPrrE 34J8 •/-0.11 PLATEAU #MB •^^Ai/^'Ar PHLOGOPTTE34.10 •/-035 PLATEAU»MB

#AKC Armstong, P., Kamp, PJJ., and Constenius, K.N., 1998 (in progress) #B Bromfield, C.S., Erickson, AJ„ Jr., Haddadin, M. A., and Mehnert, H.H., 1977, Potassium-argon ages of intrusion, extrusion and associated ore deposits, Park City mining district, UtSi; Economic Geology, v. 72, p. 837-848. #C Constenius, K.N., 1998 (this paper) #CR Crittenden, M.D., Jr., Stuckless, J.S., Kistler, R. W., and Stem, T.W., 1973, Radiometric dating of intrusive rocks in the Cottonwood area, Utah: U.S. Geological Survey Journal of Research, v. 1, p. 173- 178. #J John, D.A., Turrin, B.D., and Miller, RJ., 1997, New K/Ar and 40 Ar/39Ar ages of plutonism, hydrothem^ alteration, and mineralization in the central Wasatch Mountains,//! John, D.A., and Ballantyne, G.H., eds.. Geology and ore deposits of the Oquirrh and Wasatch Mountains, Guidebook prepared for Society of Economic Geologists Field Trip October 23 to 25,1997, Guidebook Series No. 29,p. 65-80. #K Kowallis, B.J., Ferguson, J., and Jorgensen, G.J., 1990, Uplift along the Salt Lake s^ment of the Wasatch fault from apatite and zircon fission track dating in the Little Cottonwood stock: Nuclear Tracks Radiation Measurement, v. 17, p.325-329. #M Morrissey, A. M., 1980, Element partitioning in feldspars and apatite fission track ages fix)m seven intrusive bodies of the Park City Mining District, Utah [M.S. thesis]: University of lowa. #MB Mitchell, R.H., and Bergman, S.C., 1991, Petrology of lamproites: Plenum Press, New York and London. #N Nelson, M.E, 1977, A new Oligocene faunule from northeastern Utah: Transactions Kansas Academy Science, v. 79, p. 7-13. #P Parry, W.T., and Bruhn, R.L., 1986, Pore Fluid and seismogenic characteristics of fault rock at depth on the Wasatch Fault, Utah: Journal of Geophysical Research. V. 91, p.730-744. 149

Stock range from about 5-12 Ma for apatite fission-track to 31 Ma for K/Ai hornblende; implying a protracted, possibly tiered, cooling history which began in Oligocene time. Fission-track ages of zircon and sphene and K/Ar and ^Ar/^^Ar ages of biotite and muscovite from the Little Cottonwood stock cluster around 28-30 Ma on the east-side and 24-25 Ma on the west-side. These results are compatible with exhumation of different depth levels of the pluton (ca 6 km in east, 11 km in west) and cooling from above 400° C to below 180O C, at these respective times and places. Houghton (1986) considered that this "300° thermal event" at approximately 25 Ma marked the inception of "Wasatch" (Deer Creek) fauit movements. Second, tiie Alta stock and stocks east of the Alta stock show general concordance of the low- and medium-temperature thermochronometers. Specifically, fission-track analyses of apatite, zircon and sphene 5aeld ages roughly agree with K/Ar and ^Ar/^^Ar ages of biotite and hornblende (range of closure temperatures ca. 1000-5000 C; Hodges, 1991), Comparing the fluid inclusion based paleodepths of John (1989) (ca. 1-6.5 km. Fig. 3.7) with apparent ages of the Alta, Qayton Peak, Flagstaff, Mayflower, Park Premier, Valeo, Pine Creek, Ontario stocks, indicates that these stocks experienced rapid cooling in Oligocene (ca. 36-32 Ma) time. In particular, the deepest of these stocks, the equigranular phase of the Alta (Brighton) (ca. 5.6-6.5 km) cooled from >700® C to <300o C within the uncertainty of the various U /Pb, K/Ar and fission frack thermochronometers. The degree to which this represents tectonic exhumation versus heat retention of stocks intruded at shallow depth is dependent on the assumed paleogeothermal temperature gradient which will be discussed later. 150

Study of the Comer Creek tectonite and southwestern mcirgin of the Little Cottonwood stock using K/Ar and apatite fission track age-data (i.e., no track lengtti measurements) led Evans (1985) to conclude that Wasatch fault- slip began about 17.6 +/- 0.7 Ma (K/Ar sericite alteration in fault zone) and at an average rate of 0.76 mm/yr. over the last 10 m.y. Parry and Bruhn (1986, 1987) used the ca. 17.6 K/Ar age as the time of initial movement and fluid- inclusion based paleodepth estimates of 11 km to calculate an exhimiation rate of 0.67 mm/yr. Age-data from a second apatite fission track investigation of the same part of the Little Cottonwood stock, led Kowallis et al. (1990) to conclude that exhumation of the stock associated with movement on the Wasatch fault began about 11 Ma and continued at a high rate to the present (0.68 mm/yr.) resulting in about 6.5-7.5 km of footwall exhumation; considerably less than the previous estimate. The results of ^Ar/39Ar multidiffusion experiments on K-feldspar from three different paleodepth levels of the Little Cottonwood stock show tiered age-spectra consistent with the progressively deeper crustal levels predicted by John (1989) (Figs. 3.7 and 3.18), Age comparison with U/Pb and apatite fission track ages from the same localities as the three K-feldspar samples is included in Figure 3.18 to place high and low closure-temperature boimds on the interpretation of these data. For example, the K-Feldspar sample from Silver Lake cirque appears to have excess Ar based on age comparison with U/Pb zircon data. To partially rectify cooling information lost in the higher temperature steps of the Silver Lake sample, whole rock ages from a nearby aplite dike and the Silver Lake mylonite were added as a proxy (Figs. 3.16 and 3.18). The overlapping age spectra of the two whole rocks 151

Figtire 3.18. Summary of ^Ar/^^Ar and apatite fission track data firom different paleodepth levels of the Little Cottonwood stock. Emplacement age of stock defined by U/Pb result. Paleodepth levels based on fluid-inclusion and mineralogical evidence fi:om John (1989) and Parry and Bruhn (1986,1987). Apatite fission track data from Evans et al., 1987, Kowallis et al., 1990, and P.J.J. Kamp and P. A. Armstrong (written commun., 1997). Argon spectra plotted using MacArgon 5.11 program courtesy G.S. Lister and S.L. Baldwin, Monash Univeristy, Melbourne, Australia, and University Arizona, Tucson, Arizona. Little Cottonwood Stock SILVER LAKE CIRQUE - PALEODEPTH ~ 6 km K-Feldspar (KNC62695-2) SIl Vr.R l AKi: MYI.ONIn; - PAI EODI H ll ~ b km U/reZfiton Whole Kock (Muscovite(?)| (rAV96()720-4) sn.vr:R i akf cirqui: apijtI' dikh - PAiriODiiPTH - b Wholo Rt)ck |K-l eldspar(?)| (KNC81196-3)

DRY CREEK CANYON - PALEODEPTH ~ 85 km K-Feldspar (KNC62995-2) o 20 JACOB'S LADDER - PALEODEPTH ~ 11 km K-Feldspar (KNC7195-1B)

SIl.VF R I.AKi: CIRQUI-- APAi ri l; HSSION TRACK

JACOB'S LADDliR - APAI HI' FISSION IRACK DRY CIU'IiK CANYON - APAl l I !• I ISSION TRACK

Fraction Ar Released 153

samples at the high-temperature steps (0.5-1.0 39Ar released) and the general concordance of the K-feldspar and whole rock mylonite spectra at the low- temperature steps (0.05-0.5) lead one to conclude that if there hadn't been excess argon the K-feldspar sample would have jaelded a plateau age for the higher temperature steps of about 28 Ma. K/Ar and ^Ar/^^Ar biotite ages from this part of stock are also about 28 Ma as is the zircon fission track age from the same sample. However, the fission track apatite age is from this sample is much yoxmger, ca. 12.3 +/-1.5 Ma (Fig. 3.16, Table 2). We conclude from the Silver Lake age-data that the Little Cottonwood stock experienced rapid initial cooling from a molten state >850^ C at ca. 30.5 Ma to <200® C at ca. 28 Ma, but that the pluton remained at temperatures greater than annealing temperature of apatite (ca. 105) until about 12 Ma. Age-spectra from the Dry Creek Canyon sample (ca. 8.5 km paleodepth) reflect mainly partial loss of argon consistent with either slow cooling over a protracted time interval (10-12 m.y.) and/or the effects of a post-26 Ma thermal pulse, that followed initially rapid ca. 26-27 Ma cooling. The apparent apatite fission track age of this sample is 4.9 +/- 0.8 Ma (P.J.J. Kamp and P. A. Armstrong, written commim., 1997). Argon data from the Jacob's Ladder sample can be interpreted as evidence of a ca. 17 Ma cooling episode that closed the higher temperature domains, followed by a period of slow cooling until about 7-8 Ma when the stock was rapidly cooled. This interpretation would suggest that exhimiation of the stock was episodic and that it is not justifiable to extrapolate exhumation rates for the last 7-10 m.y. back to 17 Ma or 25 Ma (e.g., Evans, 1985). Alternatively, the part of the spectrum from 0.3 to 1.0 39Ar release fractions may entirely reflect partial loss 154

related to a thermal pulse, leaving only the lower fractions as reliable indicators of the cooling history. Importantly, combining published apatite fission track based exhtimation rates and ages (Evans, 1985; Kowallis et al., 1990) with the improved temporal resolution of the lower release spectrum of the K-feldspar argon data, yields a clear definition of the timing and rate of slip on the Wasatch fault. These data rndicctte that the latest episode of Wasatch fault-slip initiated 7-8 Ma on this fatdt segment and at an average rate of 0.68 mm/yr. resulting in about 4.8-5.4 km of footwall exhumation. Reconstruction of P-T-t paths for the Little Cottonwood stock and 1-D thermal modeling of conductive heat loss of the pluton showed that it would be difficult to distinguish denudation from magmatic cooling if the geothermal gradient at the time of intrusion was 30° C/km, a rate advocated by Parry and Bruhn (1987). However, if the paleogeothermal gradient was 45° C/km or greater, a clear distinction can be made between the two processes and there would be evidence of significant late Paleogene-early Neogene exhumation of the Deer Creek footwall. Recall that displacement of the "Eocene unconformity" accounts for about 7 km of late Paleogene-early Neogene throw across the Deer Creek fault. Comparison of structured elevations of 9-10 Ma apatite age sites versus elevations of the same age for zircon tission track analyses (Kowallis et al., 1990) suggest that the Neogene palegeothermal temperature gradient was greater than 500C/km. Similar comparison of time-temperature curves modeled from K-feldspar argon diffusion data from the three depth levels of the pluton (Silver Lake, Dry Creek Canyon and Jacob's Ladder) yield a ca. 45° C/km paleogeodiermal temperature gradient. Study of present-day heat flow on the east-end of the 155

core complex lead Moran (1991) to conclude that the near-surface heat flow varied from 25-65° C/km. Collectively, these geochronologic analyses were designed to address the questions of when, and at what rate was the footwall of the Deer Creek fault tectonically breached? Unfortunately, even with the large amoimt of data available no definitive statement can be made regarding questions of denudation because of the difficulty in differentiating magmatic cooling from tectonic exhiunation which ultimately rests on assimied paleogeothermal temperature gradients. Several lines of evidence, however, can be used to argue that heat flow in the Cottonwood core complex was elevated at the time of intrusion of the Little Cottonwood stock (>450C/km) and may have remained high to the present-day. If so, the geochronologic data record significant late Pedeogene-early Miocene structurally induced cooling of the core complex. Based on these geochronologic data combined with structural evidence of Deer Creek detachment and Wasatch normal faulting we propose that footwall exhumation was partitioned by complimentary mechanisms and at different times. The Little Cottonwood stock and its coimtry rock, the core of the Cottonwood complex, were tectonically denuded initially in Oligocene and early Miocene time during collapse of the C-N thrust sheet. Continued exhumation linked to Wasatch normal faulting along the western margin of the complex, a product of Basin and Range extension, began later (ca. 7-8 Ma) and continues to the present-day. 156

3.6 COTTONWOOD METAMORPHIC CORE COMPLEX 3.6.1 Flexural-Isostatic Exhumation & Crustal Welt Metamorphic core complexes of the North American Cordillera are domal or anticlinal exposures of metamorphosed and mylonitized-brecciated rocks exhumed from mid-crustal depths beneath regional detachment faults (Jackson, 1997; Coney, 1980). We consider the footwall of the Deer Creek detachment fault, the area formerly known as the Cottonwood arch to be a metamorphic core complex because it possesses not only an arched infrastructure but also has structural domains characteristic of core complexes (Coney, 1980; Wernicke, 1992): (1) uiunetamorphosed cover terrane or superstructure that sustained a considerable amount of extension and stratal attenuation (e.g.. Box Elder Peak anticline, Tibbie half-graben. Silver Lake detachment), (2) footwall metamorphic tectonite or decollement zone, a sharp fault surface very visible in topography characterized by development of fault-rocks (breccia, microbreccia, cataclasite and myloiute) and extensive alteration (e.g.. Comer Creek tectonite. Silver Lake mylonite, marble tectonite at Silver Lake), and (3) arched basement terrane of metamorphic-plutonic- sedimentary rocks that often contain synextensional granitic bodies and metamorphic assemblages suggestive of deptiis greater than 10 km (e.g., WIB, Big Cottonwood metasedimentary rocks). We next explore crustal thickening as a cause of crystalline basement uplift and propose that buoyant rise of a crustal welt, tectonically denuded at shallow levels by detachment faulting and softened by heat and fluids derived from mafic intrusions and crustal melts, drove construction of the Cottonwood metamorphic core complex. To test this hypothesis we 157

constructed a Bouguer gravity model across the Qieyenne belt just east of the Cottonw^ood complex.

3.6^ Bouguer Gravity Model & Discussion We used 2-D Bouguer gravity modeling within the constraints provided by reflection seismic profiles, borehole data, and crustal thickness and velocity-structure results (Willden, 1965; Braile et al., 1974; Smith and Bruhn, 1984; Hall and Chase, 1989), to probe the subsurface crustal structure of the Cheyenne belt and to establish the crustal thickness beneath the Uinta Mountains uplift-Cottonwood arch. Our goal was to verify whether a relict of a formerly large Laraniide crustal welt exists. We selected this line of profile which is east of tiie plunging nose of the Cottonwood arch-core complex and west of the Uinta Moimtains for the following reasons: (1) this area has a higher likelihood of preserving the crustal welt since it lies east of where Deer Creek-Wasatch normal faulting stripped the upper crust allowing the crustal welt to buoyantly rise, and, of where there was significant assimilation of the lower crust that lead to WIB melts, (2) the surface and shallow-crustal subsurface structure is well defined by published maps and cross sections, borehole studies, and seismic reflection data (Lamerson, 1982; Bruhn et al., 1983, 1986; Bradley and Bruhn, 1988; Bryant, 1990; Bryant, 1992; Constenius and Mueller, 1996; Constenius, unpublished seismic reflection data), (3) surface elevations and relief and corresponding terrain correction affects are relatively subdued along this transect in contrast to the rugged alpine topography and large terrain affects foimd in the core complex, and (4) 3-D gravity perturbations caused by low-density Neogene basin-fill flanking the 158

Wasatch fault are minimal along X-X' whereas they cloud deep-crustal gravity signatures to the west. Observed Bouguer gravity values obtained from Cook et al. (1989) show essentially no anomaly across the Uinta Moimtains uplift which is odd because there is more than 10 km of structural relief across this inverted Proterozoic rift structure. Similar structural relief in the central Uinta Mountains produces Bouguer gravity anomalies of 50-60 mgal (Fig. 3.19) (Smith and Cook, 1985; Cook et. al., 1989). Density values used in the model were obtained by integrating published sources (Smithson, 1971; Hall and Chase, 1989), converting seismic refraction velocities to density tising formulas of Christiar\sen and Mooney (1995), and borehole logs (Constenius, 1998). The values of specific gravity used in tiie 2-D Bouguer gravity model are displayed in Figure 3.19 and the justification for their use will be discussed in Constenius (1998). The most critical rock density to define was the Uinta Motmtain Group because this >10 km thick body of rock directly overlies the Cheyenne belt. Iterative Bouguer modeling showed that the negative anomaly contributed by a supposed crustal welt could also be accounted for by low density Uinta Moimtain Group (<2560 kg/m3). Neutron formation density, sonic and velocity logs from nine wells that penetrate the Uinta Mountain Group meeisured or calculate bulk densities of 2560-2670 kg/m^ (2620 kg/m^ average). Modeling calculations were made using a modified version of the Talwani 2-dimensional modeling program (Talwani et al., 1959; C. Chase written commimication, 1997). The density structure shown in Figure 3.19 was contrasted against a reference column from the Wyoming province used by Hall and Chase (1989). X X' South Calculated Bouguer Gravity North

5 -180 Observed Bouguer Gravity -200 OL ^ gs -180 .• ...... ^ ,, , O -240 Po CO Uinta allochthon Charleston-Neho Uinta Mountains Hogsback & Absaroka allochthon Reference uplift thrust sheets Column ^•2500 I p-2540 I,.. i-p'2520 UfptrCmt V-266Q p-2670 M4mOwt 0-2870 2800 / \ \ s province UanoffCrmi Wyomms province p-2970 vVs\Vs''v ('Pr0ter020/c crust) (Archean crust) ////•////////'P r.»\ \\ \s\\s\\\sss\s\ss\\s\\\\\ / \M .1. ./ •/ // // // /•//•//• Mantle Mono P'3300 N \ Mantle p - 3300 Crustal ivelt 100 200 DISTANCE (km)

Figure 3.19. Bouguer gravity model X-X' of the Qieyeme belt showing relict Laramide crustal welt developed at convergence zone between crustal provinces. Shallow crustal structure based on interpretation of seismic reflection and borehole data for tlie Qiarleston-Nebo allochthon (Coiistenius, 1996), and published reports for the Uinta Mountains uplift, Absaroka thrust, and Hogsback thrust ^amerson, 1982; Bradley and Brulm, 1988; Bryant, 1992; Royse, 1993). Observed Bouguer gravity data from Cook et al., (1989). 160

Bouguer gravity data modeled cdong X-X' indicate that downward deflection of the Moho by a crustal welt is a permissible solution. The crustal- density structure depicted in Figure 3.19 is consistent with our interpretation that a crustal welt produced by north-south contraction across a rejuvenated Qieyeime belt crustal suture exists today beneath the west-end of the Uinta Moimtains and to various degrees beneath the Cottonwood core complex. The modeled thickness of the welt is about 5 km measured against the thickness of imdeformed Proterozoic crust to the south. If higher density values for the Uinta Moimtain Group (e.g., 2650 kg/m^) are used the thickness of the welt increases by 2-3 km. Inherent in the gravity model is a structure produced by two modes of Campanian-middle Eocene Laramide shortening that acted simultaneously to eastward-translate and invert the Proterozoic Uinta aulacogen. It has recently been hypothesized that the Uinta Mountains uplift- Cottonwood arch is an allochthonous element of the Sevier orogenic belt and that this structure is the centred member of three overlapping thrust sheets that collectively describe the Uinta allochthon (Constenius and Mueller, 1996). The Uinta BMB fault and the North Flank fault, blind thrusts that sole at depth to detach the Uinta Mountains uplift from its substrate, link with the C-N and Hogsback sole thrusts to form a crustal scale duplex (Figs. 3.2, 3.3, and 3.19; Bradley and Bruhn, 1988; Constenius and Mueller, 1996). The three thrust sheets combined to form the Uinta allochthon, a thrust-fold structure that was transported 15-20 km east in Campanian-middle Eocene time. Simultaneous with thrust-slip, the Uinta Mountains-Cottonwood arch was uplifted and folded by north-south convergence between the Wyoming and 161

Yavapai provinces. The mode of structural accommodation varied latercdly along the Uinta Moimtains with 15-30 km of north-south shortening (Stone, 1993; Constenius and Mueller, 1996; Constenius 1998) taken-up by deformation in ihe lower crust and/or channeled to shallower crustal levels. In otiier words, shortening across the Qieyenne belt was balanced by different degrees of overthrusting or imderthrusting of Proterozoic Yavapai crust against rigid cratonic rocks of the Archean Wyoming province. On the west-end of the Uinta Mountains the Uinta aulacogen has been structurally inverted yet there is no evidence that the boimding contractile faults reach the surface (i.e., blind thrusts) implying that north-south contraction here was accommodated by shortening in the lower crust (Figs. 3.19 and 3.20). In the center of the range, where there is no suggestion of a crustal welt in the gravity data and as much as 18 km and 11 km of displacement on the North Flank and Uinta BMB faults, respectively, crustal shortening was accomplished entirely by faults that cut upward toward tiie earth's surface (Gries, 1983; Qement, 1983; Smith and Cook, 1985; Constenius, 1998). Eastward from the central Uinta Mountains, displacement on the range boimding thrusts decreases and there is a correspondingly greater amoxmt of contraction in the lower crust. Refraction and gravity data suggest that a crustal welt formed in the eastern Uinta Moimtains but that the lack of detachment faulting has suppressed any tendency to rise. Summarizing, Bouguer gravity modeling indicates that a crustal welt exists beneath the Cottonwood arch and newly conceptualized structural models of the Uinta allochthon and Cottonwood metamorphic core complex link the formation of ttiis welt to contractile deformation across a 162 rejuvenated Proterozoic arustal suture followed by flexural-isostatic exhumation during detachment faulting (Fig. 3.20). The Cottonwood core complex possesses three traits common to core complexes in the Cordillera: (1) extension along a low-angle detachment fault(s) remove large sections of upper crustal rocks, resulting in isostatic uplift of tihe footwall, (2) buoyant rise of crust overtiiickened during episodes of shortening (Coney and Harms, 1984, Spencer and Reynolds, 1990; Myers and Beck, 1994) (3) a spatial and temporal link between core-complex formation and plutonism (Lister and Baldwin, 1993 and references therein). We consider the spatial association of the WIB with the Deer Creek detachment no coincidence, because it was here that the crustal welt was tectonically stripped allowing decompression melting of the crust. Our preferred model for tfie origin of WIB magmatism based on the chemical distinctiveness of the individual stocks is partial melting of calc-aUcaline mafic to intermediate crust as a consequence of decompression simultaneous witii nortiiwest-southeast oriented crustal extension and intrusion of mafic magma into the lower crust (Vogel et al., 1997). The extra heat provided by mafic intrusions combined with, decompression, would cause relatively rapid crustal melting. These magmas most likely resulted from melting of Proterozoic, calc-alkaline crust, which then could have ascended along extensional accommodation zones in the crust. The effects of WIB plutonism, thermal weakening inducing ductile behavior and inflation of magma-chambers, were important to tiie formation of the core complex (Lister and Baldwin, 1993). But, the main process responsible for creation ofthe Cottonwood core complex is isostatic uplift driven by the Cheyenne belt 163

Figure 3.20. Schematic diagram illustrating role of Cheyenne belt crustal welt in tectonic development of late Paleogene-Recent magmatism, extension and flexwal-isotatic exhumation of Cottonwood metamorphic core complex. COTTONWOOD ARCH COTTONWOOD CORE COMPLEX CC (MIDDLE EOCENE) rC CC (PRESENT) CC South ,T- North South North Utnta Mountains uplift Deer Creek Little Cottonwood Charleston-Nebo allochthon / Absaroka thrust detachment stock fault t mm.

-20" ^ h ^ i'J 5 ;• "i i •'> EWilfffliifll

Mono Mono dikes & sills / derived from PROTEROZOIC ARCHEAN crustal melts CRUST CRUST ~11 km exhumation CRUSTAL WELT related to mafic dikes & sills flexural-isostatic in loiver crust compensation ~30 km N-S crustal shortening 165

crustal welt stripped of its upper crustal load. Perhaps the most compelling argument for considering the Cottonwood area as a metamorphic core complex is that crystalline basement rocks of the Proterozoic Little Willow Formation at the core of the structure were exhumed from depths of 11 km or more in late Eocene-Recent time whereas short distances to the north and south of the Cheyenne belt cratonic crystcdline basement is gently west-dipping, undeformed, and resides at depths in excess of 11 km (Figs. 32,3.3 and 3.19; Coogan, 1992; Yoi\kee et al., 1997).

3.7 SIGNinCANCE OF LATE PALEOGENE EXTENSION In the context of regional Cordilleran space-time patterns of crustal shortening, extension, and magmatism, the Cottonwood metamorphic core complex and C-N allochthon are situated in an area in which there were profoimd changes in the style of tectonism over a ~ 10 m.y. period beginning in early middle Eocene time (ca. 50 Ma) (Constenius, 1996). Thrusting ended and extension began at about 50 Ma in the southern-most part of the Wyoming province , whereas a short distance to the south in the Yavapai province, results from Constenius (19%) and this study document that, the switch from C-N thrust and Uinta uplift related shortening to collapse and synextensional magmatism began at about 40 Ma. Thus, from early middle Eocene to late Eocene time (ca. 50-40 Ma), terrains imdergoing markedly different structural evolution (i.e., active thrusting in the south versus thrust sheet collapse ia the north) were juxtaposed across the Cheyeime belt crustal boundary. 166

The late Eocene - early Miocene record of collapse and coeval magmatism, which is preserved in half-grabens filled with volcanogenic sedimentary rocks, was previously unrecognized in this area. Past studies of central Utah concluded that extension began in Miocene time (ca. 25 Ma or 17 Ma) (Royse, 1983; Hopkins and Bruhn, 1983; Gallager, 1985; Parry and Bruhn, 1987) and considered assemblages of volcaniclastic rock to be erosional valley fill (e.g., Runyon, 1977), In fact, this region has experienced two distinct phases of extension: (1) late Paleogene collapse of Sevier belt fold-thrust structures that was coeval witti development of a metamorphic core complex (ca. 40-20 Ma), followed by (2) Neogene Basin and Range extension (ca. 17-0) (Constenius, 1995, 1996, 1997).

Recognition of late Paleogene extension is important because contemporjiry geologic models of the Wasatch Mountains suggest that the Wasatch normal fault has sustained more than 11 km of Neogene exhumation and cast this as t5rpical of Basin & Range fault systems (Parry and Bruhn, 1986; Parry and Bruhn, 1987; John, 1989). In contrast, studies based on topographic relief, depth of basin-fill and stratigraphic projection of the Wasatch fault placed the amount of vertical offset in the range of 1.5 -4.6 km (e.g., Zoback, 1983). Paleodepth studies of the depth of emplacement of plutonic rocks and formation of fault-rocks, based mainly on fluid-inclusion, fault-rock and petrographic observations of Wasatch footwall rocks, advocate extremely high rates of Neogene denudation (Parry and Bruhn, 1987). However, the timing and genesis of footwall exhimiation and tectonic interpretations based on these factors were poorly constrained due to the unrecognized earlier episode of late Paleogene crustal extension that was of 167

equal or greater magnitude. We propose that a significant component of the crustal exhimiation and the formation of the deep-seated fault-rocks and half-grabens were created by the Deer Creek detachment-fault system, a progenitor to the Wasatch fault. Furthermore, we contend that the 11 km of Deer Creek- Wasatch footwall exhumation documented from the Cottonwood metamorphic core complex is unique to this locality (Salt Lake-Provo segments of the Wasatch fault zone) and atypical of the magnitude of extension associated with Basin and Range normeil faults regionally and for other segments of the Wasatch fault (refer to Schwartz and Coppersmitii, 1984). Rather, convergence of tectonic elements at this juncture, which includes a crustal welt constructed during the Laramide contractional episode, resulted in generation of crustally derived melts and exhimiation driven by flexural-isostatic compensation of the Cottonwood arch, WIB, and C-N sole fault. The pronounced magnitude of crustal rebound is related to tectonic unroofing of the welt by the Deer Creek and Wasatch normal faults, and extensional reactivation of the crustal boundary created magmatic accommodation zones for magmas of the Wasatch igneous belt.

3.8 CONCLUSIONS

(1) The Charleston, Nebo, and Uinta BMB thrust are structural elements of the Santaquin culmination, a crustal scale duplex that accommodated more than 100 km of shortening. Progressive unconformities in Coniandan-Campardan synorogenic strata (ca. 90-84 Ma) record the main 168 phase of C-N thrusting (the upper horses of the duplex) that involved ~80 km of shortening. Thrusts of the C-N allochthon were subsequently folded and in some cases reactivated to accommodate slip at the base of the Santaquin culmination on the Uinta BMB fault. Tectonic transport of the lower horse of the Santaquin antiformal duplex (~20 km), took place from Campanian through late Eocene time (ca, 82-40 Ma) based on sjmorogenic growth strata deformed by fold-thrust structures.

(2) Growth of the Santaquin culmination lead to interned imbrication of the frontal C-N allochthon by east-dipping backthrusts. Renewal of Nebo thrust shortening in the opposite direction produced unusual yotinger over older thrust relationships mapped as the "Deer Creek thrust." The east- dipping structural fabric of the C-N hanging wall played a critical role in the structural accommodation of late Cretaceous-early Paleogene internal contraction of the C-N allochthon, a role that would later be repeated in an opposite sense when these inclined detachment surfaces facilitated collapse of the allochthon.

(3) Late Eocene-middle Miocene (ca 40-16 Ma) backslippage of the Santaquin thrust-wedge, the structure that drove internal contraction of the C-N allochthon, caused the sense of slip on inclined thrust planes to be reversed and the creation of unusual half-grabens. These half-grabens have bowl-shaped map patterns and profiles, exhibit stratal growth geometries in the synextensional basin-fill on both sides of the basin, and the basin- bounding faults are guided by the same mechanically weak and broken 169

Stratigraphy that had accommodated crustal shortening. Unusual bedding geometries in the grabens which heretofore were explained by models involving diapirism and depositioiuil onlap, are rollover anticlines and synextensional stratal growth formed by bed rotation above listric normal faults.

(4) Structural relationships observed on the northern margin of the C- N allochthon adjacent the Cottonwood core complex suggest that beginiung in late Eocene or early Oligocene time (ca. >36.4 Ma) the C-N sole thrust was extensionally reactivated and its hanging wall was transported west- southwest on the Deer Creek detachment fault. The initiation of extei\sion coincided with the inception of local and regional magmatism and formation of the low-angle detachment faults and metamorphic core complex in the hinterland of the Cordillera (Armstrong and Ward, 1991; Dickinson, 1991; Axen et al., 1993; Constenius, 1996). Movement on the detachment continued through Oligocene-early Miocene time coeval with emplacement of WIB intrusive rocks, displacing the Eocene unconformity at least 7 km in the western part of the Tibbie half-graben. Formation of phyllonite and cataclasite observed in the Comer Creek tectonite are the product of west- southwest subhorizontal sheeir beneath the 5-10 km thick C-N allochtition.

(5) A large component of the present-day structural relief and tiie 11 km of exhimiation recorded in the footwall of the Deer Creek fault were the product of a late Paleogene-early Neogene extensional episode. This interpretation contrasts with earlier studies that attribute the present Wasatch 170 front landscape and total structural relief across the fault to be solely the product of Basin and Range normal faulting (ca. 25- or ca. 17- 0 Ma). Models of Basin and Range normal faulting that have incorporated large normal offsets (-11 km) may have overestimated, by a factor of two or three, the actual magnitude of extension characteristic of Basin and Range fault systems. 171

CHAPTER 4 TECTONIC EVOLUTION OF THE JURASSIC-CRETACEOUS GREAT VALLEY FOREARC, CALIFORNLA.: IMPLICATIONS FOR THE FRANOSCAN THRUST-WEDGE HYPOTHESIS 4.1 ABSTRACT Interpretation of seismic reflection data and restoration of depositional geometries of Cretaceous forearc-basin strata in the northwest Great Valley of California provide important controls on structural reconstructions of the western margin of the Sacramento Valley and northern Coast Ranges. Monoclinal eastward dips of Great Valley Group strata and fault systems striking northwest-southeast are features proposed as evidence for a west- dipping blind Great Valley-Frandscan sole thrust and related backthrusts, but instead are expressions of bedding geometry that resulted from folding of the Paskenta and related synsedimentary normal faults, depositional onlap, and a major structural-stratigraphic discontinuity. The discontinuity separates east- dipping, Aptian and yoimger Great Valley Group strata from beds of lower Great Valley Group and Coast Range ophiolite tiiat were deformed and erosionally or structurally truncated by mid-Cretaceous time. Dip divergence imaged between the supracrop and subcrop of the discontinuity is not imique to the ancient Great Valley forearc but is also observed in modem forearc basins. Advocates of the Franciscan thrust-wedge model have also proposed that west-dipping, shingled patterns of seismic events imaged beneath the Sacramento Valley are imbricate thrust slices of Great Valley Group. This hypothesis, however, is incompatible with borehole, potential-field and seismic-refraction data that characterize the Sacramento Valley basement as ophiolitic. Seaward-dipping reflections in the ophiolitic basement of the ni

Sacramento Valley are analogous to layering developed in the oceanic crust of volcanic rift margins or generated along mid-ocean ridges. Thus, late-stage diastrophic mechanisms are not required to interpret a forearc that owes much of its present-day bedding architecture to processes coeval with deposition. Thickening of the Great Valley Group stratigraphic section (Valanginian-Turonian) in the hanging walls of the Paskenta, Elder Creek, and Cold Fork fault zones, combined with attenuation or complete omission of pre-extensional units (including the Coast Range ophiolite) and geometric evidence based on seismic reconstructions suggest that these faults are Jurassic-Cretaceous normal faults that developed in a submarine setting. Down-structure views of the Great Valley outcrop belt simplify otherwise complex map relations and portray the Paskenta and related faults as half- graben botmding faults which accommodated sigiuficant northwestward tectonic transport of hanging wall rocks. Significantly, these faults sole into the Coast Range fault, an enigmatic forearc structure that juxtaposes rocks of Franciscan complex (blueschists) with rocks of the Coast Range ophiolite and Great Valley Group that have sustained only zeolite-grade metamorphism. Discovery of Jurassic-Cretaceous crustal-scale extension in the Great Valley forearc suggests that a significant part of Coast Range fault-related attenuation (~ 15 km) developed early in the history of the subduction complex.

4.2 INTRODUCTION

4.2.1 General Exposed in the California Coast Ranges is the classic analog of a subduction-type convergent margin from which sedimentary and structured 173 prcxiesses in forearc basins and acaetionary prisms have been modeled (e.g., Dickinson and Seely, 1979). This Mesozoic-Paleogene Coast Range subduction assemblage is composed of the Franciscan complex, the Coast Range ophiolite and the Great Valley Group. Although outcrop relationships generally are complicated by rapid lithofades variations, polyphase deformational histories and burial by late Neogene-Recent sediments, decades of carefxil geologic mapping and interpretation have led to the development of an extensive base of information on the stratigraphy, paleontology and structure of these units. Nevertheless, studies have been imable to fully reconcile stratigraphic and structural observations (e.g., Qen, 1992; Ingersoll and Dickinson, 1992; also see discussion below). For example, there still is no consensiis on the nature of the contact between the Coast Range ophiolite and the Franciscan complex, the genesis of the Paskenta fault zone and other faults that cut the Great Valley Group, and the accommodation of shortening by inferred "backthrusts" with no apparent surface traces. This paper uses new subsurface data combined with published surface mapping and biostratigraphy to define the three-dimensional structural and stratigraphic architecture of the northwest Great Valley. Using insights gleaned from this dataset we propose hypotheses that resolve ttie apparent disparity between structural- versus stratigraphic-based interpretations of Coast Range subduction complex tectonics.

4J2J2 Significance of Problem Study of California convergent-margin tectonics recently has focused on two prevailing models: 1) Late Cretaceous-Paleocene extensional 174 iinroofing and exhumation of the Franciscan complex (Harms et al., 1992; Krueger and Jones, 1989; Jayko et al., 1987; Piatt, 1986) and 2) Maastrichtian- Recent Franciscan thrust wedging above blind thrusts along the Coast Ranges- Great Valley boundary (Wentworth et al, 1984; Namson and Davis, 1988; Wentworth and Zoback, 1989; Unruh and Moores, 1992; Unruh et. al., 1995; Wakabayashi and Unruh, 1995). Juxtaposition of blueschist and greenschist rocks has been dted as evidence that the Coast Range fault is extensioneil (Harms et al., 1992; iCrueger and Jones, 1989; Jayko et al., 1987; Piatt, 1986). Alternatively, Ring and Brandon (1994) used kinematic data from shear zones found in proximity to the Coast Range fault to argue that it is a contractional feature. The Coast Range fault is interpreted variously as (Fig. 4.1): 1) the roof thrust of a Franciscan thrust wedge (Wentworth et al., 1984; Godfrey et al., 1997), 2) a Tertiary high-angle, east directed reverse faidt (Jayko and Blake, 1986), 3) a low-angle normal fault related to late Cretaceous and early Tertiary extensional unroofing of the Franciscan accretionary complex (e.g., Hanns et al., 1992) and 4) a west-directed thrust (e.g., Dickinson and Seely, 1979). These models have been applied regionally and predict that rocks of the Franciscan complex. Coast Range ophiolite and Great Valley Group have experienced phases of compressional, extensional and strike-slip faulting. Seismic reflection, refraction and potential-field data have added a new dimension to arguments about stratigraphic and structural evolution of this important convergent margin, and have led to concepts of intercutaneous thrust wedging of the Franciscan complex beneath rocks of the Great Valley Group (Wentworth et al., 1984; Wentworth and Zoback, 1989; Glen, 1990; 175

Figure 4.1 Montage of tectonic models of the Great Valley forearc basin and Coast Range subduction complex. Rock unit abbreviations: CRO, Jurassic Coast Range ophiolite; EUMW, extended upper mantie wedge; FCX, Cretaceous Franciscan complex; CVS, Upper Jurassic-Cretaceous Great Valley Group; SA, Sierran arc; UM, upper mantle. Modified from Brandon and Ring, 1994. 176

WEST TECTONIC MODELS EAST Great Valley forearc model - Dickinson & Seely, 1979 > CRO

10 km Franciscan thrust-wedge model - Wentworth et al., 1984 CRO ^ GVS

Franciscan thrust-wedge model - Wakabayashi & Unruh, 1995

m N\SSN\SV \ \>

EUMW^

Extensional model - Harms, et al., 1992

Thin-skinned thrust model - Suppe, 1978 177

Unruh et al., 1991, Unruh and Moores, 1992; Unruh et. al., 1995) and variations of this concept (Meltzer, 1988). These studies advocate the hypothesis ttiat eastward-tapering wedge(s) composed of Franciscan complex. Coast Range ophiolite and Great Valley Group rocks, have indented eastward, delaminating Great Valley strata. The wedge is inferred to be bounded below by a west-dipping blind thrust, and above by an east-dipping roof thrust ("backthrust") or series of east-dipping backthrusts, creating 'triangle zone" structures similar to features described in the Cordilleran foreland fold and thrust belt (Jones, 1982, 1996; Price, 1986). The magnitude of horizontal shortening depicted in various thrust-wedge models ranges from 30 to 75 km (Wakabayashi and Unruh 1995; Unruh et. al., 1995; Jachens et. al., 1995). Interpretation of the effects of tiiese possible backthrusts in the Sacramento basin poses an interesting dilemma because most are shown as bedding parallel faults in shale units with no dip divergence between the hanging wall and footwall. The seismic interpretation by Wentworth et al. (1984) shows no backthrusts; Uru^ and Moores (1992) show a single backthrust with limited displacement; and Unruh et al. (1995) show the backthrust detachment horizon as a mid-Cretaceous unconformity. Outcrop data show no major offsets or repetition of biostratigraphic markers and ramp-flat thrust relations are not identified on reflection seismic data. Yet, backthrusting is required by the wedge model. Wentworth (1984), recognizing the problem posed by the absence of major repetition of Great Valley stratigraphy, suggested that the imbricate wedges are progressively older downward and westward within the Great Valley Group; prior to deformation the sequence exhibited eastward onlap, thus permitting older. 178 western sections to be thrust beneath younger, eastern sections. Because outcrop data show no major offsets of biostratigraphic markers, backthrusts in the section must be parallel to bedding, acting as slip surfaces to accommodate wedge-related shortening (i.e., delamination). Bedding-parallel faults with no discernible stratigraphic separation cannot thicken the section because there is no stratal repetition. However, the Paskenta, Elder Creek, Sunflower Flat and Cold Fork fault zones, which have been interpreted as backthrusts by Qen (1990), cut across bedding and juxtapose rocks of different ages. A basic limitation of studies advocating the Franciscan thrust-wedge concept has been the lack of a spatially extensive network of seismic reflection and borehole data in the key areas of contact between the Franciscan complex. Coast Range ophiolite and the Great Valley Group at the western margin of the Great Valley. Consequently, outcrop relationships west of the limits of their seismic profiles were projected into the subsurface. Furthermore, these interpretations assimie that significant variations in bedding attitude seen in outcrop are structural discontinuities related to backthrusts above a tectonic wedge or imconformities related to growth of the wedge (Unruh et. al., 1995). The tectonic significance of diese outcrop features is problematical, however, because some bedding discontinuities once considered to be roof thrusts or progressive unconformities of the Franciscan wedge have been identified as buttress unconformities on the margins of large submarine canyons (Nilsen and Imperato, 1990; Williams, 1997). In previous studies, many critical structural relationships in the northern Coast Ranges-Great Valley region were not seismically imaged in the subsurface; as a result, the thrust-wedge models and other models were largely unconstrained in time and space. 179

We are able here to properly interpret the intricate outcrop and subcrop relations by using borehole data and an extensive grid of 2-D seismic profiles that spans the Great Valley outcrop. A novel and recurring theme of our investigation has been to differentiate between stratal architecture that reflects contractile deformation versus forearc depositional geometries and synsedimentary faulting.

4^3 Geologic Setting The Franciscan accretionary complex, Cocist liange ophiolite and Great Vcdley forearc-basin sequence in northern California have distinctive rock types and structural styles, but are closely linked tectonically as components of a subduction-tj^e convergent margin (Fig. 4.2). The Franciscan complex has been interpreted as a late Mesozoic and Cenozoic accretionary prism because of its deformational style, high-pressure/low-temperature (blueschist) metamorphism, the presence of entrained slivers of ophiolite and marine sediments, and its paleogeographic setting oceanward of the Great Valley forearc basin and the Sierran magmatic arc (Hamilton, 1969; Dickinson, 1971). The Franciscan complex is overlain tectonically by the Coast Range ophiolite and Great Valley forearc-basin strata. The Coast Range ophiolite is a dismembered remnant of Middle-Upper Jurassic (ca. 170 to 155-150 Ma) oceanic crust that consists mainly of serpentinite, gabbro, and basaltic volcanic rocks (Piatt, 1986; Robertson, 1990, Dickinson et al., 1996). Ophiolitic rocks are depositionally overlain by marine strata of the Great Valley Group, and are juxtaposed across the Coast Range 180

Figure 4.2. Index map showing relationship of study area to elements of Great Valley and Coast Ranges. Axis of Great Valley magnetic and gravity higji shown as dashed line (Cady, 1975). 181

FRANCISCAN COMPLEX 122 "n

220" N COAST STUDY RANGE FAULT AREA

Redding

Red Bluff COAST u^Chico RANGE Paskenta OPHIOLTTE B ^ \ V \ Sutter Buttes GREAT V \ \ < VALLEY Sacramento GROUP

San Francisco

Scale 100 km 182 fault with the Franciscan complex (Godfrey et al., 1997). A striking feature of the Coast Range fault is that, although it is contractional in that it places older rocks of die Coast Range ophiolite over younger Franciscan rocks, geobarometric analysis of Franciscan rocks (26 to 40 km paleodepths) suggests that it cuts out a vast (~ 15 km) structural thickness to juxtapose zeolite and blueschist fades rocks (Harms et al., 1992 and references therein). Rocks of the Upper Jurassic-Cretaceous Great Valley Group, exposed in a 250-km-long strip along the west side of the Sacramento Valley, form a stratigraphic succession more than 15 km thick. These rocks form an eastward-tapering sedimentary prism that rests nonconformably on disrupted fragments of the Coast Range ophiolite in the west and onlap Sierran basement to the east (Chuber, 1961, Safonov, 1962, Brown and Rich, 1967, Robertson, 1990, Dickinson et.al., 1996). The Great Valley Group, which consists primarily of mudstone intercalated with sandstone and conglomerate lenses, is made up of submarine-fan, basin-plain, slope and submarine- canyon sediments deposited in an asjmunetric, nortii-south trending forearc basin (Robertson, 1990; IngersoU and Dickinson, 1981; Brown and Rich, 1967; Suchecki, 1984; Williams et al,, 1997). The complex stratigraphic and structural relationships preserved in these three major tectonic elements record the evolution and eventual extinction of a subduction margin that persisted for more than 150 m.y. (Page and Engebretson, 1984).

4^.4 Paskenta, Elder Creek and Cold Fork Fault Zones Historically, investigation and conceptualization of Jurassic-Cretaceous convergent-margin tectonics in northern California has involved the pivotal 183 geology near Paskenta^ CA, located in the northwestern Sacramento Valley (Fig. 4.3). Numerous disparate h5rpotheses have been generated to explain the origin of an intriguing array of faults (Coast Range, Paskenta, Elder Creek and Cold Fork faults) exposed in this area and to relate these faults to convergent- margin tectonics. The Paskenta, Elder Creek and Cold Fork fault zones are enigmatic structures that strike southeastward across the Great Valley outcrop belt, and are characterized by marked stratal omission and radical stratal thickness changes. Northward, the faults attenuate the Coast Range ophiolite, sole into the Coast Range fault, and do not offset the Franciscan complex. Based on fades relationships, shoreline reconstructions showing 100 km of left-lateral offset, and the present steep attitude of the faults, Jones and Irwin (1971) proposed that the faults were large left-lateral strike-slip faults. Noting the low angle between the fault plane and bedding IngersoU and Dickinson (1981) suggested that these faults were syndepositional thrusts. Wentworth et al., (1984) inferred that the faults are hybrid left-lateral tear faults that bend soulliward, parallel to bedding, to become roof thrusts of tectonic wedges. However, Wentworth et al. (1984) comment that this relatioiiship is either buried under Cenozoic cover, or is imrecognized, and is not shown on their profiles. Kinematic indicators and structural data interpreted as showing westward tectonic transport on the Elder Creek fault zone have been used to argue that these faults are east-dipping thrusts (Glen,1990). A "sdssors- tectonics" thrust model proposed by Blake and Jayko (1993) also suggested that the faidts are thrusts, some with large displacements (~ 90 km). Sedimentologic, stratigraphic and structural evidence suggest that these 184

Figure 4.3. Geologic index map showing location of seismic profiles (designated by circled numbers) and wells used in this study. Geologic basemap adapted from Jennings (1977). Rock imit and fault abbreviations: CFF, Cold Fork fault zone; CRO, Jurassic Coast Range ophiolite and serpentinite-matrix melange; CRF, Coast Range fault; EC, Elder Creek fault zone; FCX, Cretaceous Franciscan complex; Ju, Upper Jurassic Great Valley Group; KT, Klamath terrane; Kl, Lower Cretaceous Great Valley Group; Ku, Upper Cretaceous Great Valley Group; PK, Paskenta fault zone; Qa, Quaternary alluvium; SS, Sulphur Springs fault; Tt, late Tertiary Tehama and Red Bluff formations; and Tv, Tertiary volcanic cover . Abbreviated well names: AHA, American Himter Alvarez #1; GMH, Gulf Merle Hamm #1; HC, Himible Capital #B-1; HF, Humble Fredericksen #1; HM, Hiunble Michael #1; and, SV, Shell Vilche #2-5. Location of seismic line 51 of Wentworth et al. (1984) shown as 51 (dashed line). 185

122 30'W 122 W 1

\ N N N \ S \ \ N V \ V*Y^

Red Bluff

40° N / 40° N AHA Paskenta MH i

51 J Sites anticline 3SP30'N Figure 10 3Sf30'N

10 km

122 30'W 122" W 186 faults are characterized by stratigraphic omission and expansion, cut bedding at low angles, were coeval with Jurassic-Cretaceous sedimentation, and have apparent left-lateral offset of bedding (Jones et al., 1969; Moxon, 1990; Sucheki, 1984; Vogel, 1985). Suchecki (1984) considered the Paskenta fault to be a synsedimentary normal fault but considered the Elder Creek and Cold Fork faults to be related to westward thrust movement associated with underthrusting of the Klamath terrain. Moxon (1990) concluded that all three faults were synsedimentary listric normal faults, but that the Paskenta fault also experienced oblique reverse-fault reactivation (based on interpreted footwall drag), following its inception as a left-lateral oblique extensional fault. Detailed analysis of small-scale fault structures and folds, and outcrop- scale field relations along the Paskenta fault zone were used by Vogel (1985) and Vogel and Qoos (1985) to conclude that the fault zone was a down-to-the- north zone of synsedimentary normal faulting that unroofed the rising Franciscan accretionary prism. Regional observations used by Vogel (1985) to support his hypothesis were: older beds dip more steeply than yoimger beds, time-equivalent imits are coarser and thicker on the north sides (hanging walls) of the major faults as compared to the souttiem sides, and older formations show greater offset than yoimger formations. These problematic fault zones provide critical insights into the evolution of the Great Valley forearc basin and the three-dimensional architecture of a fossil subduction system. They also permit evaluation of structural models for the evolution of convergent margins. We provide new constraints that limit possible models for the evolution of the forearc basin by integrating information from stuiidal studies with seismic-based 187

reconstructions of the depositional setting of the Great Valley Group, subsurface mapping of the Paskenta and Elder Creek fault zones, and by constructing a downplimge projection of the northernmost part of the Great Valley monocline.

4.3 STRATAL RECONSTRUCTIONS OF THE GREAT VALLEY FOREARC AND PASKENTA FAULT ZONE

4.3.1 Description of Data & Method

Through use of proprietary seismic and well data, we compiled a 600- km grid of deep, high-quality, seismic reflection profiles, and complete log suites from five deep wells (including dipmeter, biostratigraphic and velocity- log data) from the Paskenta area (Fig, 4.3). Two vintages of data were used in this study: 1) 48-channel, 24-fold, 32-8 Hz vibroseis data with maximum source-to-receiver offset of 1707 m acquired in 1976-1977 by Texaco Exploration and Production, Inc., 2) 120-channel, 30- or 60-fold, 12-57 Hz vibroseis data with maximtim source-to-receiver offsets ranging from 2783 to 4291 m acquired in 1981-1983 by Shell Western Exploration and Production Company. The source-receiver o^set is an important acquisition parameter because the more recent lines, which have longer source-receiver offsets, were able to image strata dipping as steeply as 60^ or more in contrast to the older profiles which appear to only resolve strata dipping less than 45-50®. The high-quality industry profiles were reprocessed using ProMAX software, measurably improving image quality and extending record lengths to 8 seconds. Geoquest workstation software was used to facilitate correlation, interpretation and manipulation of the data. Since much of our investigation involved resolving the nature of structural or stratigraphic 188 terminations and corresponding dip divergences seen on the seismic data (e.g., sedimentary onlap and shingling versus low-angle faulting), the ability to flatten numerous seismic events on the workstation enabled us to remove structural overprint and illuminate more subtie dip relations and thus to develop credible argimients regarding the genesis of observed reflection patterns. Synthetic seismograms were constructed using GMA software to establish formation and dipmeter ties from the deep boreholes to the seismic profiles. Stratal reconstructions based on the seismic reflection data were constructed as follows. First, time-migrated profiles were converted to depth so that basin structure could be represented with no vertical exaggeration. Depth-conversion velocities for Great Valley Group rocks were derived from borehole velocity surveys and sonic logs from deep wells (Fig. 4.4). We had no direct measure of the rock velocity of the ophiolitic basement so we relied on the seismic refraction results of Godfrey et al. (1997) which showed that the velocity of ophiolitic basement beneath the Sacramento Valley ranges from 5.5-6.5 km/s at shallow levels and increases to 8.1 km at the base. Four stratigraphically based velocity regions were defined on each profile for the depth conversion: Upper Jurassic Coast Range ophiolite (6.0 km/s), Jurassic- Cretaceous Stony Creek succession (4.2 km/s). Cretaceous Great Valley Group (3.5 km/s), and Cenozoic deposits (2.0 km/s) (Fig. 4.5). Four migrated and depth-converted seismic profiles are shown in Figure 4.6 with color coded interpretations of the rock packages (Fig. 4.5). Following depth conversion the profiles were incrementally horizon flattened to remove structural overprinting in two steps. In this procedure. 189

Figxire 4.4. Graph of interval velocity versus depth for Great Valley Group rocks from borehole velocity surveys and sonic logs. Interval velocities used in depth conversion of seismic lines TX-3 and TX-5 are indicated. Solid symbols are interval velocity data from Cretaceous Great Valley rock package and open symbols are interval velocity picks from underlying Jurassic- Cretaceous Stony Creek rock package. 190

0

Cretaceous Great Valley Depth Conversion Velocity (3.5 km/s) 2 ^

Stony Creek Depth \ * • -o, ^ Conversion Velocity (4.2 km/s) Si ^ **-dL. ca Gulf Merle Hamm #1 4 Q Humble Michael #2^ • 0 Shell Vilche §2-5 J • Humble Fredericksen §1 \ *o American Hunter Alvarez #1 X A i JU I . I Jii 2 3 4 VELOCITY (km/s) 191

Figure 4.5. Chronostratigraphy of the Great Valley monocline, northwest Sacramento Valley highlighting stratigraphic age-relations associated with lower Cretaceous discontinuity imaged on seismic data, penetrated by wells, and, seen on regional maps and satellite imagry of outcrop belt. Note that no major unconformities pimctuate the Great Valley Group witii exception of submarine unconformity related to indsion of Williams Canyon (Williams et al., 1997). Abbreviations: CRO, Coast Range ophiolite; CZ, Cenozoic rock package; E, epoch; P, rock packages defined in figure 5; Pf, petrofades of Ingersoll (1976); and, Z, benthonic foraminiferal zonation (Almgren, 1986). Sources for stratigraphic column; Bertucd (1983), Ingersoll (1979), Robertson (1990), Gradstein et al. (1994), Dickinson et al. (1996), Williams and Graham (1997). 192

STRATIGRAPHY Tehama Fm loreno Sh illllll Maastrichtian Mokelumne Fm PimnSi Starky & Wiaters Ss Sacramento Sh Campanian Kione Fm Forbes Fm Dobbins Sh Santonian illiams CynmT GiundaSs Funks Sh Coniancian Sites Ss Yolo Sh Ttironian riskeCkSh Cenomanian SaltCk Cgl

Lodoga Albian Fm

Aptian III DISCO NTINU. Bairemian Hauterivian Valangiman Bemasian

Tithoman Kimmeridgian CRO volcanopelagic Oxfordian Callovian Bathoman Coast Range Ophiolite 193

Figure 4.6. Montage of seismic profiles from the northwestern Sacramento basin that image the Pakenta and Elder Creek fault zones. Cretaceous discontinuity, and ophiolitic basement. Shown in A-C and D-F are stratal reconstructions of the Great Valley forearc basin which restore the basin to its Cretaceous configuration £is imaged on lines TX-3 and TX-5, respectively. The sections are displayed as true scale depth sections. Oblique and strike-oriented images of the Paskenta and Elder Creek faults and the ophiolitic basement are displayed in G and H, respectively. Five color coded rock packages are delineated on the seismic profiles: (1) Cenzoic (white) - meiinly Pliocene rocks of the Tehama Formation that onlap truncated, east-dipping Cretaceous Great Valley strata; (2) Cretaceous Great Valley (yellow) - monoclinal, east- dipping panel of Aptian-Santonian strata that onlap or are in fault contact with Stony Creek Formation or ophiolitic basement; (3) Jurassic-Cretaceous Stony Creek (green) - westward thickening prism of sedimentary rocks bounded by discontinuity at top and conformable with Coast Range ophiolite at base; (4) Jurassic Coast Range ophiolite (purple) - basement imit that tapers from 7.5-9.0 km thick imder the Sacramento basin to 2 km or less in or under the Great Valley outcrop belt; and, (5) Jurassic-Cretaceous Franciscan complex (yellow-green) - tectonic contact with Coast Range ophiolite. Blue line in B and E is flattened horizon sxuface. Refer to Figure 4.5 for chronostratigraphic relationships of rock packages. Location of "fork structure" designated by letter "F". OEPTH (km) DEPTH (km) DEPTH (km) DEPTH (km)

DEPTH (km) DEPTH (km) DEITH (km) DEPTH (km) 195

following Steno's principle of original horizontality, we assumed that seismic events corresponding to stratal surfaces used as flattening horizons were horizontal to gently dipping at the time of deposition. The discontinuity between the Upper Jurassic-Lower Cretaceous Stony Creek Formation and overlying Cretaceous Great Valley Group rocks was flattened first, followed by flattening of a younger Upper Cretaceous unit in the Cortina petrofades. Because parts of the seismic events that were flattened are erosionally tnmcated, parts of the restored horizons project above the present surface. In all cases, we chose to extend the projected horizons along curved paths that parallel the surface near the western ends of the profiles. Whereas we recognize that other horizon-projection trajectories could be chosen, we iteratively selected ones tfiat resulted in geologically credible depictions and minimal distortion of the data. To this end, we established a set of examples from modem forearc basins as a standard of comparison for our horizon- flattened sections (Fig. 4.7; Williams and Graham, 1997). Lastly, we constructed "paleoseimic lines"; a creative data enhancement method in which stratigraphic and paleontologic-based water-depth information (Dickinson and others, 1987) was integrated with the horizon flattening of time-migrated seismic data. This tjrpe of paleobathymetric reconstruction allowed us to compeire seismic time-section examples from modem forearc basins with our computer generated images of the ancient Great Valley forearc (Fig, 4.7, lower left).

4.3.2 Interpretation of Reflection Seismic Profiles The most intriguing reflection geometries seen in seismic lines TX 3, 5, 196

Figure 4.7. Comparative examples of active Cenozoic forearc basins imaged on seismic reflection profiles and hypothetical reconstruction of the Qretaceous Great Valley forearc basin. Modem sections were chosen to highlight basin morphologies resulting from extensionally driven subsidence, and submarine sedimentation and erosion. Notice shingling of strata in detached basins, dip divergences between forearc strata and substrate in all excimples ("fork structures"), unconformities within forearc basin assemblages (e.g.. West Luzon Trough and Hesperides basin-B), drape of strata across bathymetric steps, and dramatic changes in stratal geometry in different parts of the Hesperides basin. These forearc attributes are captured in paleoseismic reconstruction of the Great Valley forearc basin using seismic line TX-3 combined with water depth information from Dickinson et al. (1987). Bathymetric step stylized on TX-3 to account for dramatic arcward thinning of basin-fill (compare with West Luzon Trough and South Shetland forearc). All sections shown in two-way travel time and vertical exaggeration of 5:1. Open-ellipse on left of examples is ca. 3000 m depth mark reference. Strata shown in modem forearcs are of late Paleogene to Neogene age (no data are available for Arica basin). Sources for seismic: Coulboum and Moberly (1977), Coulboum (1981), Moore et al. (1982), Hayes and Lewis (1984), Dobson et al. (1991), and, Maldonado et al. (1994). Figure adapted partly from Williams and Graham (1997). 197

MODERN FOREARCS Trenchward Arcward

ARICA BIGHT FOREARC, CHILE SOUTH SHETLAND FOREARC, ANTARCTICA Arica basin llesperides basin - A

Detachment

WEST LUZON TROUGH, PHILLIPINES Detachment (?)

SOUTH SHETLAND FOREARC, ANTARCTICA Hesperidrs basin - B

Detachment

ALEUTIAN FOREARC, ALASKA Atkin basin

SOUTH SHETLAND FOREARC, ANTARCTICA Onlap Trench Hesperides forearc basin - A s.l. Accretionary prism

ANCIENT GREAT VALLEY FOREARC, CALIFORNIA Lower Crctaceons NW Sacramento basin Trench fill S-0 s ^ ^

Uceanic crust

SO km

Detachment SUNDA FOREARC, CENTRAL JAVA, INDONESIA

Trench Accretionary prism

4 s 3 km

ceanic crust

SCALE 50 km VE - 6:1

10 km DIP 198

12 and SH 808 are the pronounced dip divergences between steeply east- dipping events in the Gretaceous Great Valley rock package and west-dipping events in the Jurassic-Cretaceous Stony Creek and Coast Range ophiolite (CRO) units, and the high-amplitude events tiiat demarcate the boundaries of the Coast Range ophiolite (Fig. 4.6a, 4.6d, 4.6g & 4.6h). The steep eastward dip that characterizes the Cretaceous Great Valley rock package is merely a basinward continuation of the Great Valley monocline, whereas beneath these dipping units the westward thickening Jurassic-Cretaceous Stony Creek rock package forms a discordant sedimentary prism. In contrast, the CRO is a west-tapering imit that thins from 7.5 to 9 km in the east to less than 2 km in the west (a relcttion previously imrecognized). The CRO abruptly thins concomitant with rapid westward thickening of the Stony Creek rock package giving rise to a fork-shaped pattern of reflectivity (*F' in Figures 4.6a-f). The 'fork structure" is so-named because of the convergence of reflections at the fork produced by west-dipping events in the CRO, the Stony Creek-CRO contact, west-dipping events in tiie Stony Creek, tfie Stony Creek-Cretaceous discontinuity, and east-dipping events in the Cretaceous Great Valley. The discontinuity surface that separates the Cretaceous Great Valley and Stony Creek-CRO rock packages is interpreted to be tfie product of both normal faulting coeval with and a hiatus in, forearc sedimentation (Figs. 4.6c and 4.6f). Where the Stony Creek-Cretaceous dip discordance is stongest, the discontinmty is interpreted as a fault zone ttiat separates strata of the Stony Creek package in the footwall from Cretaceous Great Valley strata in the hanging wall. Eastward, beyond the limits of tiie normal fault, the discontinuity is a condensed section or angular imconformity 199

(nonconformity on CRO) developed on CRO and Stony Creek subcrop and onlapped by rocks of the Cretaceous Great Valley rock package. The normal fault surface blends seemlessly with the unconformity because both geologic elements formed in a submarine setting. The age of the discontinuity and the angularity of bedding across the surface are well defined on seismic line TX-3 by including paleontologic, petrofades and dipmeter data from the American Hunter Alvarez #1 well. Here, the discontinuity has been folded into a gentle anticline, and the borehole penetrates the discontinuity at the crest of the anticline. Within the borehole, the supracrop is composed of Aptian-Albian rocks of the Lodoga Formation (J-2 benthonic foraminiferal zone; Fig. 4.5) dipping 25-270 east, whereas the Hauterivian (?) subcrop of the Stony Creek Formation (K beniiionic foraminiferal zone) dips 2-9° west. Thus tiiere is about 30° of dip- divergence across the discontinuity. It is noteworthy that Aptian-Albian rocks of the Lodoga Formation are much thinner in the American Himter #1 well (520 m) than to the west (>2500 m). The hiatus associated with the discontinuity is more difficult to gauge because of the resolution of the biostratigraphy and xmcertainties in the age designations that were assigned, but could be as much as 10 m.y. if lower Aptian and Barremian rocks are truly absent. The youngest rocks drilled in this well are Tiironian (H benthonic foraminiferal zone, 100-1570 m) near the surface and the oldest foimd at the base of the well are Tithonian (M bentiionic foraminiferal zone, 3764-4285 m). Linkage of the discontinuity with the system of large faults that characterize ttie Paskenta area «ire seen on seismic profiles TX-5 and SH-808. These profiles, along with TX-12, also provide evidence that the Paskenta 200

fault does not involve basement beneath the Sacramento Valley. As was the case with line TX-3, these profiles image the impressive dip divergence of subcrop and supracrop reflections across the discontinuity but, in addition, the discontinuity can be followed to the surface where it intersects the surface trace of the Paskenta fault. Along with the Great Valley monocline, the discontinuity has been folded such that it dips steeply northeast (-50°) necir the surface and gradually flattens with depth, and to the northeast, as the effects of Neogene deformation diminish. Borehole information from the Gulf Merle Hamm #1 is less clear but, based on seismic correlation and limited biostratigraphy, the discontinuity was picked at a depth of about 2127 m. The supracrop associated with the discontinuity in this case has been described as Albian or younger(?) and the subcrop interpreted as pre-Albian. At the site of the borehole penetration of the discontinuity there is little angular variation in bedding aaoss the surface. Rocks above and below the discontinuity dip 10-20O northeast. However, to the south of the well, seismic line SH-808 images 20-60O of dip divergence across the discontinuity (Fig. 4.6g). Collectively, these seismic profiles offer evidence that the discontinuity is an angidar unconformity and/or normal fault zone (since it attenuates section) that separates the Cretaceous Great Valley and Jurassic-Cretaceous Stony Creek rock packages. Stratal reconstructions that remove the effects of post-Cretaceous deformation and restore the forearc basin to roughly its Cretaceous configuration provide insight into the processes that resulted in a surface that was the product of both nondeposition/erosion and normal faulting (Fig. 4.6b, 4.6c, 4.6e, & 4.6f). The geometry of reflections in the Cretaceous Great 201

Valley (CGV) rock package have the appearance of strata in a submarine half- graben. That is, CGV units in the restored profile are folded into an asymmetrical syncline in which units in the southwest limb dip 30° east and gradually flatten upsection, and units in the northeast limb of the syncline dip gently 0-5® west. All of the CGV units terminate against the discontinuity; western imits are inclined and abut the discontinuity at moderate angles whereas, to the east, the units are subhorizontal and onlap the discontinuity (Fig. 4.6c & 4.6f). Normal fatdting in the early stages of CGV sedimentation resulted in bed rotation and high-angle terminations against the discontinuity (fault) and creation of a bathymetric depression. Slip on the normal-fault system eventually ceased, and the depression was filled by continued deep-water sedimentation. Therefore, the origin of the discontinuity surface and the hiatus associated with it, are dependent on its position in the forearc basin. On the southwest end of the profiles and in the adjoining Great Valley outcrop belt, the discontinuity is a normal fatdt (i.e., Paskenta fault), whereas to the east, such as where the CGV onlaps the CRO, the surface is a nonconformity. The discontinuity is color coded in Figures 4.6c and 4.6f to reflect where the discontinuity owes its origin to normal faulting (red) or to nondeposition/erosion followed by depositional onlap (yellow). Examples from modem forearc basins emphasize that the genesis of dip divergence, detachments, unconformities, and significant variations in basin geometry along strike are ordinary elements of forearc basins. Common to all of the examples is the development of dip divergence between the forearc fill and the basin substrate. Dip divergence that developed across the 202 discontinuity that separates the fill from its substrate is the product of detachments with normal sense of offset (e.g., Arica basin), depositional onlap of bathymetric depressions (e.g., Atkin basin), or a combination of both (e.g.. West Luzon Trough). Unconformities within the forearc fill are evident in all of the examples. The examples from the South Shetland forearc highlight the radical along-strike variation in forearc structure (compare Hesperides basins A & B). In our paleobathymetric reconstruction of the ancient Great Valley forearc (Fig. 4.7, lower left), aspects of the Arica basin and the Hesperides basin - A are evident: detachment-related bed rotation, landward thinning and onlap, and a bathjmietric-structural step. The bathymetric- structural step shown in Figure 4.7 was incorporated into the ancient Great Valley reconstruction to accoimt for the anomalously thin, or complete lack of, Jurassic-lower Cretaceous Stony Greek-Lodoga strata beneath the central- eastern Sacramento VeiUey (i.e., area of sedimentary bypass and/or submarine erosion). Further evidence of a bathjonetric step associated with the Paskenta fault is foimd in submarine mass movement deposits described by Vogel (1985) which are stmunarized in the next section. The Hesperides basin - B with its fork-structure geometry and lack of detachment serves as an analog for the ancient Great Valley south of the Paskenta area. South of the Paskenta area, the origin of the discontinuity is solely related to east-directed depositional onlap onto beveled rocks of the Stony Creek Formation and ttie Coast Range ophiolite. However, in the Paskenta area, interpretation of the restored versions of seismic profiles TX 3 & 5 and surface mapping indicates that the Paskenta and related fault systems were active from Tithonian-Valangian througjK Turonian time. Following 203 cessation of slip on the normal fault at about 90 Ma the mode of deep-water sedimentation changed from detachment-dominated deposition to foundering of the basin and encroachment on the Sierran arc. Late Cretaceous sedimentary rocks onlapped and prograded eastward across the Coast Range ophiolite at this latitude in the basin, thereby forming the discontinuity (nonconformity) imaged on the east-ends of these profiles. Before this time, deposition of Great Valley Group rocks had been restricted to west of the "fork-structure", that is, west of the point where the CRO begins to taper to the west. In contrast, elsewhere in the Sacramento basin, this fimdamental shift in the depositional pattern of the basin began in Aptian time (ca. 120 Ma). The pattern of regional onlap of Cenomanian and yoimger Great Valley Group strata has long been established (e.g., Chuber, 1961). What is newly defined here is the initiation of basinwide onlap of ophiolitic basement and the Sierran arc in Aptian time and the recognition that a major discontinuity (unconformity) partitions the Great Valley Group. The diagrammatic representation of the discontinuity in Figure 4.4 shows that there is little if any hiatus associated with the discontinuity in the Great Valley outcrop belt to the west, but that the hiatus is more than 60 m.y. to the east beneath the Sacramento Valley. The existence of a Lower Cretaceous discontinuity was postulated by Peterson (1967) who noted significant erosional stripping of the subcrop (~ 450-600 m) and reworked fossils in coarse clastic rocks overlying the erosion surface. This hypothesis lacked initial widespread acceptance because there is littie or no hiatus associated with the discontinuity in the Great Valley outcrop belt. Shortiy thereafter, however. 204

Dickinson and Rich (1972) and later Ingersoli (1983) recognized changes in sandstone petrology across the surface. The minimal hiatus associated with the discontinuity in the outcrop belt, in rocks which comprise the western part of the fork structure, stems from the fact that this area was the Tithonian to Barremian depocenter of the forearc basin. There is however a significant structural manifestation of the discontinuity that appears on geologic maps at the base of the Lodoga Formation (Brown and Rich, 1%1; Rich, 1971; Rich unpublished mapping) and is clearly discernible in airphotos and satellite imagery. Composited Landsat images show the undulating pattern created by low-amplitude east- plunging folds in the Stony Creek Formation that terminate against the discontinuity, versus overlying beds of the Lodoga Formation that form linear ridges paralleling the discontinuity (Fig. 4.8). We believe that structural and petrofades changes across the discontinuity combined with tiie basinwide change in depositional pattern imply that the forearc was subjected to a change in tectonic regime beginning in Aptian time. It is no coincidence that accelerated Farallon-North American plate convergence rates and/ or shallowing of the subducting oceanic plate in the Aptian, mechanisms hypothesized to accoimt for the Sevier orogeny (backarc convergence), would involve a change in depositional-structural regime in the foreeirc (Burchfiel et al., 1992). Specifically, dynamic subsidence (e.g., Gumis, 1992) associated with flattening of the subducted Farallon plate (Currie, 1997) may have led to the inimdation and eventual burial of the western part of the Sierran arc; the record of this episode is the regional pattern of depositional onlap on basement rocks of the 205

Figure 4.8. Montage of Landsat imagery of Great Valley outcrop belt, NW Sacramento basin which highlights strike changes in bedding attitude across the Qretaceous discontinuity. Exposed west of the discontinuity are subcrop strata of the Stony Creek Formation that have been folded into low amplitude folds that terminate at the discontinuity (prominent truncations marked T), whereas supracrop rocks of Lodoga-Riunsey formations parallel the discontinuity. Discontinuity merges with Paskenta fault to north. 206 207

Sacramento Valley.

4.4 DOWN-STRUCTURE VIEW OF GREAT VALLEY OUTCROP BELT

Recognition of the sjmsedimentary nature of the Paskenta, Elder Creek, Cold Fork and related fault systems suggests that these structures could be better examined from the perspective of a down-structure view, which removes the monoclinal tilt of Great Valley strata (Mackin, 1950), rather tiian by considering map-plane relations alone. The down-plxmge projection of the northern part of the Great Vcdley monocline, which includes the enigmatic Paskenta and other fault zones, is shown in Figure 4.9. This figure is a composite of three viewing angles, 30° E, 45° E, and 60° E, an integration that was required to match the variation in bedding attitude of the outcrop belt and minimize distortion. The chronostratigraphic map of Moxon (1988) served as a basis for the overall bedrock geology of the monocline, and maps by Chubar (1961), Dondanville (1958), Jayko and Blake (1986), Jones and Bailey (1973) and Murphy et al. (1969) were used to add map details. It is apparent on the resultant crustal-scale "cross-section" (Fig. 4.9) that the Paskenta, Elder Creek, Cold Fork, Sulphur Springs and Oak Flat faults (and adjoining unnamed faults) collectively form a large half-graben, boimded by a nortii-stepping system of three listric normal faults to the south and two antithetic faults to the north. The horizontal dimension of the composite structure is about 50 km. Attributes of extensional faults are seen in the profoimd stratal thickening of synextensional units in the hanging walls, particularly those of Valanginian-Barremian age, coupled with 208

Figure 4.9. Downplunge view of Great Valley outcrop belt, NW Sacramento basin, segregated into biostratigraphic zones. Great Valley Group subcrop beneath Tehama Formation shown in grey, as established from seismic correlation and borehole data. Paleontologic age of subcropping units from wells labeled (black) next to well symbol. Diagram constructed from chronostratigraphic map of Moxon (1988) combined with detailed map relations fovmd in Chubar (1961), Dondanville (1958), Murphy et al. (1969), Jayko and Blake (1986), and Jones and Bailey (1973). Abbreviations of wells, faults and localities: AHA, American Hunter Alvarez #1; GMH, Gulf Merle Hamm #1; HF, Humble Fredericksen #1; HM, Humble Michael #1; and, SV, Shell Vilche #2-5; SFFZ, Sunflower Flat fault zone; and YBJ, Yolla Bolly jimction. Polee to Faults Fold Axes HF Paskenta fault zone Campanian JL HM ... .. V Elder Creek fault zone Santoniun Sites anttclme -6- 6* Fruto synclin

iurvman Cold Fork fault zone GMH

Cretaceous (7 Sulphur Springs fault

rri wwi Coast Range Fault Aptiaii Quntemary-late Tertiary Discontinuity SofUpninn (Wifliams Cyn) 1^3 Blttterivinn-B^netnian Franciscan Cpniancian Vahngitti(itt-$eTTia^ian complex Turonian TithoHian-KimmeridgiaH ^Klamath terrane JO km Cenomanian CpMSt Range opltiplite

N) ovO 210 omission of footwall rock units. The amount of apparent displacement on the fault systems measured on Figure 4.9 totals about 22 km. Stratigraphic thickening and truncation across individual faults is as high as 5.5 km and 6.0 km (i.e., Valanginian-Barremian strata involved in Elder Creek fault and Cold Fork fault zones, respectively). Tracing the Paskenta, Elder Creek and Cold Fork faults down section shows that they merge and together have thinned and cut out much of the Stony Creek Formation (Valanginian-Kimmeridgian section) and completely removed tmits of the Coast Range ophiolite. These faults sole into the Coast Range faidt at or near the YoUa Bolly Junction (Jayko and Blake, 1986), at which point hanging wall rocks of the Klamath terrain are juxtaposed with footwall rocks of the Franciscan complex. Relict slivers of the Klamath terrane, dismembered by normal faults that collectively have about 8 km of apparent displacement, are also present at the YoUa Bolly Junction. The yoimgest rocks involved in the normal faulting are Turonian strata cut by the Sulphiu: Springs fault and overlying Coniadan imits that display drag folding. The apparent stretching direction is north-south as exemplified by the major north-dipping listric normal faults in the plane of this section, but, as developed in the descriptions that follow, the true extension direction was to the northwest. Hence, the outcrop belt offers us a spectacular oblique slice through a Cretaceous half-graben situated within the Great Valley forearc. Structural and stratigraphic effects of the individual faults are svunmarized as follows; 1) The Paskenta fault has 8 km of offset measured at the top of the Tithonian. Net thickening of Valanginian-Barremian strata across the fault is about 2.8 21 1

km, and possibly 1 to 2 km of Kimmeridgian-Tithonian strata are trvmcated by the fault although it is difficult to estimate because of internal structure in the footwall.

2) Valanginian-Barremian rocks are displaced 7 km across the Elder Creek fault, Valanginian-Barremian strata show about 5.5 km of net stratal thickening from the Elder Creek footwall to hanging wall, and approximately 4 km of Tithonian-Berriasian rocks are omitted across the fault.

3) The linked system of synthetic faults that comprise the Cold Fork fault zone have in the range of 3-5 km of offset measured at the top of the Albian. The Sulphur Springs fault, which is an antithetic fault of the Cold Fork system, has 2.5-3.0 km of top Albian offset, stratigraphic expansion of Hauterivian-Barremian strata across the Cold Fork fault system is on the order of 5 km near the fault, and the thickness of these rocks balloons northward to exceed 10 km (Moxon, 1988). About 5.5 km of Valanginian strata are truncated by the fault system.

4) Two southeast-striking faults mapped as merging to the northwest are exposed in the Simflower Flat area (Moxon, 1988; Jones and Bailey, 1973). The upper splay is a low-angle normal fault that cuts upper and middle Valangiiuan strata, and the map pattern suggests that movements on this fault were sjnichronous with Hauterivian-Barremian sedimentation. The lower splay is an anomaly wittiin this overall extensional belt in that it duplicates middle-upper Valanginian strata. If the fault is contractile as 212 mapped, then the episode of shortening related to this fault was exceedingly brief (pre-Hauterivia to Barremian). As noted by Moxon (1988), the stratal repetition is based on only two megafossil localities in close proximity to faults, and therefore may warrant further biostratigraphic investigation. Stratal thickening and truncation values estimated from the down- structure diagram (Fig. 4.9) compare favorably with estimates by Jones and Irwin (1971) of over 3 km of Tithonian truncation associated with the Elder Creek fault and over 6 km of Valanginian termination related to faults of the Cold Fork system. These authors also noted that the character of equivalent sedimentary imits changes across these faults, with coarse-grained units found in the hanging wall block versus finer grained units in the footwall. The measured orientation of small-scale faults and axes of small to mediiun size folds foimd in the vicinity of the Paskenta fault zone were used by Vogel (1985) to conclude that the fault zone was the site of significant northwest-southeast directed subhorizontal extension from Tithonian through at least Valanginian time, that there is no evidence for post- Cretaceous slip on the fault zone, that the faults and folds have been rotated by Neogene uplift of the Coast Range, and that the orientation of tectonic transport of the Paskenta hanging wall was southeast-northwest. Importantly, he established that tfie structures were syndepositional based on evidence that faults and folds are overlapped by vounger undisturbed beds, sandstone layers have been ductilely deformed, sandstone dikes are present, slickensides and lineations are rare, and there is no evidence of cataclasis. Measurements of similar small-scale structures near the Paskenta fault zone and across the Elder Creek fault zone by Glen (1990,1992) led him to conclude 213 that these faults originated as diastrophic structures that accommodated thrust displacement to the west. While acknowledging tiie marked similarity between his study section and that of Vogel (1985), and also noting two cases of undeformed bedding that flank folded units, he advocated a diastrophic origin for the structures and dismissed Vogel's (1985) results because Vogel interpreted northwest-southeast transport and failed to differentiate fault- proximal structures from those in other parts of the Great Valley Group. The anomalous and complex nature of the Paskenta outcrops was emphasized in a rebuttal to Glen (1990) by IngersoU and Dickinson (1992), who differentiated structures foimd in the Paskenta area from those found in other parts of the Great Valley. Kinematic data collected by Vogel (1985) and Glen (1990) support oiu: hypothesis of large-scale synextensional normsd faulting coupled with west or northwest transport of the Paskenta and Elder Creek hanging wall blocks. Stereographic plots of fault-plane and fold-axis data from the Paskenta fault zone (Vogel, 1985) in which the data have been rotated to bed-horizontal to remove the effects of late-stage folding (eastward tilt of outcrop belt), are shown in Figure 4.8. Subsurface mapping, combined with the mapped surface trace of the Paskenta fault zone, shows that, in its present configuration, the Paskenta fault zone is a folded fault that dips to the east or northeast. Northwest- or west-directed displacement of the Paskenta hanging wall and other major fault blocks of the half-graben system shown in Figure 4.8 would have resulted in the transport of rocks upward and out of the Great Valley monocline as we view it today. Prior to uplift of the Coast Range, hanging wall rocks were merely displaced to iiie west, but later uplift resulted 214

in erosional stripping; this explains why complete biostratigraphic units, some 5-6 km thick, are absent from the Great Valley monocline. Lastly, down structure views of the Great Valley Group outcrop belt give the impression that Upper Jurassic and Cretaceous synsedimentary normal faulting may have been more widespread and not just confined to the northern reaches of the Great Valley forearc basin. Notice in Figure 4.8 that beneath the Valanginian-Barremian-age normal fault labeled "A", a ~30-km- long segment of the Coast Range ophiolite appears to be highly attenuated compared to thicker parts of the same unit to the north and south. Mapping by Chubar (1961), Brown and Rich (1961), Brown (1964), and Rich (1971) (also see Williams and Graham, 1997) upsection from the attenuated Coast Range ophiolite, show a nimiber of high-angle faults with surface traces oriented ENE-WSW that cut across N-S-striking exposures of Great Valley Group. Viewed down structure, most of these faults are normal faults. If these faults are Mesozoic structures, they provide an important structural reference by which to gauge the presence or absence of later bedding parallel faults that would laterally displace the trace of these faults. Significantly, no such lateral displacement of faults has been detected in this area.

4.5 COAST RANGE OPHIOLITE REFLECTIONS: EVIDENCE FOR VOLCANIC RIFTED MARGIN

Depositional processes considerably different from those responsible for forearc basin sedimentation accotmt for the imusucd seismic reflectivity imaged beneath rocks of the Great Valley Group in the western and central parts of the Sacramento Valley. Examples of sub-Great Valley Group 215 reflectivity are seen in seismic profiles SH-798 and SH-780/CO-3 (Fig. 4.10) and in Unrtih et al., 1995. The substrate beneath the Great Valley Group in this part of the Sacramento Valley has been characterized variously as Sierran- Klamath granitic basement (Wentworth and others, 1984), ophiolitic basement (Cady, 1975; Harwood and Helley, 1987; Jachens, 1995; Godfrey et al., 1997 and references therein), lower Great Valley Group tnmcated beneath a mid-Cretaceous unconformity (Dumitru, 1988), or thrust-imbricated lower Great Valley Group strata in the footwall of the Franciscan thrust-wedge and Sierran crystalline basement beneath a mid-Cretaceous angular unconformity (Wakabayashi and Unruh, 1995; Unruh and others, 1995). It is noteworthy that studies which characterized the basement as ophiolitic used borehole samples, potential fields, and, seismic refraction and reflection data, whereas those advocating Sierran basement and/or imbricated lower Great Valley Group relied solely on seismic reflection data. Integration of all sources of subsurface data including our new seismic reflection coverage of the Great Valley outcrop belt, and correlation of Coast Range ophiolite outcrop with seismic and gravity profiles suggest that the Sacramento basin is floored by the Coast Range ophioUte. Several published seismic profiles show a shingled pattern of west- dipping events between 2.5-5.0 seconds two-way fravel time that was interpreted as thrust repetition of lower Great Valley rocks (e.g., Unruh et al., 1995). In this interpretation it is axiomatic that crystalline rocks are acoustically transparent in contrast to layered and highly reflective rocks of the Great Valley Group. Unruh et al. (1995) dte interval velocities (from stacking-velocity analysis of 5.0-5.2 km/s) for the west-dipping basement 216

Figure 4.10. Patttems of seismic reflectivity from the ophiolitic basement of the Sacramento basin compared with volcanic margins of the Atlantic Ocean basin. 217

West Shell Line 798 Hme Migration

Ophiolitic Basement Sacramento Basin

Shell Line 798

Shell-Conoco Lines 780-3

Humble Fredericksen #I

Offshore Nova Scotia, Canada ^ 3 S Seaward-dipping reflections in oceanic basement 218 reflection interval as evidence of imbricated lov^er Great Valley Group rocks. Velocity estimates and coherent reflectivity by themselves, however, are not sufficient criteria to uniquely ascertziin the composition or character of the basement. Furttiermore, the interval velocities cited by Unruh et al. (1995) are not reliable because of the poor velodty-resolving ability of oil-industry data at these depths (limited dimensions of recording array; see Yilmaz, 1987). Even taken at face value, these rock velocities are much higher than borehole velocity measurements of lower Great Valley Group rocks (ca 4.2 km/s; Fig. 4.4). Alternatively, several lines of evidence lead us to conclude that rocks of the Great Valley Group rest depositionally on the Coast Range ophiolite and to propose that the unusual basement reflectivity is analogous to reflectivity seen on volcanic rift margins as seaward dipping reflections (i.e., manifestation of inter layered volcanic and volcaniclastic units). The distribution of the Coast Range ophiolite based on modeling of seismic refraction data, gravity, and magnetic data and tabulation of boreholes that penetrated ophiolitic rocks beneath the Great Valley Group in the Sacramento basin recently has been investigated and summarized by Godfrey et al. (1997). Borehole data combined witfi seismic-refraction and potential- field data from the Sacramento Valley indicate that the basement is part of the Coast Range ophiolite. In their description of the basement beneath the Sacramento Valley, Harwood and Helley (1987) state that 'from Sutter Buttes north to Chico, a number of wells penetrate basement rocks reported as diorite, gabbro, norite gabbro and serpentinite. Not surprisingly, this part of the Sacramento Valley is characterized by large positive gravity and magnetic anomalies that mark the northern part of the inferred Great Valley ophiolitic 219 basement (Cady, 1975). Some thin sections of gabbro and diorite from wells in this area, contained in a collection of thin sections of basement rocks held by the California Academy of Sciences, reveal remarkably fresh coarse- and medium-grained clinopyroxene-plagioclase rocks that show pronounced cumulate textures. These textures, previously unrecognized, provide supporting evidence for the interpretation that the source rocks for the Great Valley anomaly are oceanic crust". We add the following descriptions of rock samples from three deep wells (Fig. 4.3) in the vicinity of our seismic grid to these observations and to material compiled by Godfrey et al. (1997). Basement rocks drilled and cored from 3778 to 3888 m in the Humble H.C. Fredericksen #1 well (section 16, T20N, R3W, Glenn Coimty, CA; Fig. 4.3), have been described either as greenstones (Chuber, 1961) or altered volcanics witti recrystallized quartz vesicle fillings (Godfrey et al., 1997). Petrologically, the 3838 m (12,592 ft) part of the core has been described as a hydrothermcdly altered dadtic lapilli tuff with edipote, chlorite and pyrite throughout the rock or in fractures; post-alteration deformation in the form of horizontally oriented slickenlines in vertical shear zones was also noted (Holt, R.D., 1955, Humble Oil and Refining Company internal memo). In a summary of the petrography of cores from the Himible Capital Company #B- 1 well (section 3, T16N, RIW, Colusa Coimty, California) P.H. Masson (1953, Hxmible Oil and Refining Company internal memo) described three rock types from cores spanning a 3055 m to 3087 m (10,024-10,130 ft) depth range: actinolite greenschist (probably altered from basic igneous rock type) with matrix of finer actinolite, quartz and feldspar; gabbro, fine to mediimi grained; and, metadadte, porphyritic with plagiodase phenocrysts in a &ie-grained 220 groimd mass of equigranular quartz and feldspar, brecdated with fragments cemented by microcrystjdline silica, weak flow structure in fragments. Thin sections from sample depths of 3054 m (10,020 ft) and 3087 m(10,127 ft) have recently been described by Godfrey et al. (1997) as altered basalt with phenocrysts of pyroxene, amphibole, and feldspars, and an extremely altered, but foliated, volcanic flow. Rocks at the base of ttie Shell Vilche #2-5 well (section 5, T27N, R4W, Tehama Coimty, California) simply have been described as serpentinite on the mud log. Thus, it is clear that Great Valley Group rocks rest nonconformably on the Coast Range ophiolite in the central and western parts of the Sacramento Valley. Nevertheless, the question remains, what gives rise to the reflectivity observed within the ophiolite? We propose that the CRO reflections, based on their west-dipping, shingled geometry are the product of interlayered volcanic flows, volcaniclastic rocks and mafic intrusions similar to those beneath volcanic rifted margins. Marine seismic reflection and potential field investigations along the Atlantic margin of North America have established that seaward-dipping reflections within oceanic crust are a defining characteristic of such margins (e.g., Holbrook et al., 1994; Keen and Potter, 1995; Oh et al., 1995). A comparison of Coast Range ophiolite "seaward- dipping" reflections seen on two seismic profiles from the Sacramento Valley and an example from offshore Canada are shown in Figure 4.9. The Humble Fredericksen #1 well, which drilled 32 m of ophiolitic basement, has been projected about 2 km onto the plane of seismic section SH-780/CO-3, situated near two groups of seaward-dipping reflections. Dipping reflectivity in tiie ophiolitic basement beneath ttie Sacramento Valley could also be a byproduct 221 of accretionary structures formed within oceanic crust at spreading centers (Ranero et al.l997 and references therein).

In either case, coherent patterns of dipping reflectivity are documented in oceanic crust and serve as analogs for the Coast Range ophiolite that lies beneath the Great Valley. As was the case in the preceding discussion of Great Valley reflection geometry, stratal geometries presupposed to be of diastrophic origin can more easily be explained as the constructional patterns of "depositional" systems.

4.6 CONTINUITY OF OPHIOLITIC BASEMENT 4.6.1 Gravity Modeling - Data and Methods We used Bouguer gravity modeling within the constraints provided by improved definition of the Coast Range ophiolite seen on our reflection seismic profiles, borehole data, and the regional geophysical results of Godfrey et al. (1997) to probe the subsurface continuity of the Coast Range ophiolite and to establish the geometry of the Franciscan complex beneath the ophiolite. Our goals were to verify whether surface exposures of the Coast Range ophiolite are physiczilly separated from or contiguous with ophiolitic basement imaged beneath the Sacramento Valley and to trace structural elements mapped at the surface to deep levels in the crust.

Using seismic refraction results, earthquake travel-time studies, and modeling of potential-field data to establish crustal velocity-density structure and depth to Moho, Godfrey et al., (1997) linked ophiolite exposed in 222 the outcrop belt to basement rock beneath the Sacramento Valley. In contrast, Jachens et al. (1995) concluded from magnetic models that expostures of the Coast Range ophiolite were detached from west-dipping ophiolitic basement that underlies the Great Valley and Coast Ranges. Ruppel (1971) modeled the Coast Range ophiolite as a west-dipping slab with a density of 2870 kg/m^, Bouguer modeling was optimized by selecting a transect (A-A', Fig. 4.3) superposed on seismic profile TX-5 and parallel to a detailed gravity profile of Ruppel (1971; Line 500). The Bouguer gravity model initially used the deep crustal structure of Godfrey et al. (1997) and density data from Bailey et al. (1964), Ruppel (1971), Cady (1975) and Godfrey et al. (1997). Lateral changes in crustal thickness were patterned after seismic refraction and earthquake traveltime results from Puis and Mooney (1990) and Oppenheimer and Eaton (1984). Sources of observed Bouguer gravity data included Ruppel (1971) for the west end of the profile and regional data from Chapman et al., (1974) and Oliver et al., (1980). Density values for rocks of the Great VaUey Group and Franciscan used in the model were obtained by integrating published sources and borehole logs (Fig. 4.11). Values of median specific gravity of sandstones of the Great Valley Group and Franciscan complex published by Bailey et al. (1964) ranged from 2550-2590 kg/m^ and 2600-2650 kg/m^, respectively. The wet density of 30 field samples reported by Ruppel (1971) was 2350-2590 kg/m^ (average 2480 kg/m3) for rocks of the Cretaceous Great Valley rock package, 2470-2650 kg/m3 (average 2570 kg/m^) for rocks of the Stony Creek Formation, 2570- 2730 kg/m3 (average 2650 kg/m^) for rocks from the Franciscan complex., and 2270 to 2620 kg/m^ (average 2440 kg/m^) for serpentinized samples of 223

the Coast Range ophiolite. Neutron formation density logs from the Shell Vilche #2-5 and Gulf Merle Hamm #1 wells measured bulk densities of 2240- 2480 kg/m^ (2400 kg/m3 average) and 2420-2620 kg/m^ (2480 kg/m3 average) for rocks of the Cretaceous Great VaUey rock package, and 2450-2740 kg/ (2620 kg/m^ average) and 2440-2630 kg/m^ (2500 kg/m^ average) for rocks of the Stony Creek Formation, respectively. Deep crustal and mantle densities used in the model approximate values used by Cady (1975) and Godfrey et. al., 1997; 2950 kg/m^ for the Coast Range ophiolite, decreasing to 2770 kg/m^ at shallow levels due to increased alteration; 3250 kg/m^ for the Coast Range ophiolite mantle (Great Valley ophiolite mantle of Godfrey et. al., 1997), 2820 kg/m3 for the lower crust of the Sierran magmatic arc, and 3300 kg/m^ for the upper mantle. Modeling calculations were made using a modified version of the Talwani 2-dimensional modeling program (Talwani et al., 1959; C. Chcise written communication, 1997).

4.6^ Gravity Modeling Results The Bouguer gravity model is consistent with our interpretation that the Coast Rcm.ge ophiolite floors the Sacramento Valley and that thiimer, dismembered parts of the same ophiolitic unit are exposed at the surface (Fig. 4.11). Beneath the Sacramento Valley the CRO is modeled as a 9 km-thick body that overlies a 6 km-thick remnant of CRO mantle; both of these oceanic imits were obducted onto the Sierra arc in the Jurassic eind now overlie lower crustal rocks of the Sierran arc (see Godfrey et al., 1997). West of the "fork structure" the CRO thins dramatically to 3.0-3.5 km under the Great Valley outcrop belt, and instead of overlying CRO mantle and Sierrein arc rocks, the 224

Figure 4.10. Bouguer gravity model A-A' of the Great VaUey monocline based on regional seismic refraction and gravity results of Godfrey (1997), detailed Bouguer gravity models (Ruppel, 1971), interpretation of seismic reflection profile (TX-5), borehole data from the Gulf Merle Hamm #1 (GMH), and surface mapping by Maxwell (1974). Location of "fork structure" designated by letter "F". A wsw «) -30 > -40 6 -50 Calculated Bouguer Gravity ^ -60 (J -70 Observed Bouguer Gravity

-80 o -90

§ -100 Sacramento Valley Coast Range Seismic profile TX-5 Fault GMH Great Valley forearc basin p-2600 77 p-2250 (A"-400) (A = - 500) // /I - 2770 (A = +120) p • 2470 (A = - mO) '///A . P-28S0 p - 2S50 (A = -100) '{A = +200; p - 2570 (A = -HO) Coast Range ophwlite p - 2950 (A= +300)

CRO mantle Franciscan comvlex - 3250 (A= +350) 27u5 (A= -195} KsW Sierran arc ^^\'p.2820<&^-M) /^/^\\\\\\\\\\ Moho p - 2710 X' P • 2820 fA = - 480) (A = -590)^?\^^^ / / / f / / / s s s Mantle / s \ s \ \ n.3300

DISTANCE (km) 226

CRO rests on rocks of tiie Franciscan complex. Tracing the CRO to the surface from beneath the outcrop belt, the CRO thins to about 1.0-1.5 km-thick, is modeled as becoming increasing serpentinized (i.e., less dense and lower velocity) toward the surface, and has been folded such that it is overturned and dips steeply to the west. Overall, the configuration of modeled structural elements bears a striking resemblance to the long-standing structural model of the Great Valley forearc presoited by Dickinson and Seely (1979). However, we elaborate on their model and propose a modified sequence of events regarding the Coast Range faidt and its associated metamorphic gap. This model also places important limitations cm the Franciscan thrust-wedge model because tiie CRO is not cut by a hypothetical sole thrust and the CRO does not extend 50 km or more to the west as modeled by Jachens et al., 1995. Furthermore, because we have reinterpreted supposed "backtiirusts* as syndepositional normal faults and stratal onlaps, and have observed no significant thrusts or backthrusts cutting Great Valley Group strata, we conclude tiiat the wedge model is a permissible mechanism only if restricted to crustal material beneath the CRO. Lastly, we conclude that a significant componoit of Coast Range fatdt- related attenuation of the rock column is related to a phase of crustal extension concurrent with Jurassic-Cretaceous sedimentation. Evidence of this tectonic episode are first seen in the abrupt tiunning of the CRO west of the 'fork structure" and the corresponding onlap and westward thickening of the Stony Creek Formation. We infer from these relations that the present configuration of the CRO does not reflect its original thickness or that of any 227 obducted CRO mantle that may have been present beneath it. Rather, elements of the CRO west of the "fork structure" are tectonically thinned remnants of ttie CRO xmit ttiat was originally 73-9.0 km thick. Slip on a normal fault system that was a precursor to faults of the Paskenta system created a tectonic moat that accommodated deposition of over 6-7 km of Jurassic-Lower Cretaceous Stony Creek Formation. As tiiis episode of forearc extension evolved, the structural level of normal faulting locally stepped-up to the top of the Stony Creek and faults of the Paskenta system developed. Recall that faults of the Paskenta, Elder Creek and Cold Fork fault systems sole into the Coast Range fault and tiiat there is significant stratal attenuation associated with these fault systems. A comparison of the thicknesses of undeformed parts of the CRO and CRO mantle in the subsurface of the Sacramento Valley with the tectonically thinned vestiges of the CRO in the outcrop belt to tiie west, suggests that the thickness of ophiolitic rocks removed by Jurassic-Cretaceous normal faulting may range as high as 8 to 14 km. The magnitude of CRO attenuation observed here is similar to estimates of the amoimt of rock colimm removed (ca. 15 km) along the Coast Range fault in tiie Diablo Range, 200 km to the south (Harms et al., 1992).

4.7 CONCLUSIONS 1) The Paskenta, Elder Creek, Cold Fork and related fault zones are a network of normal faults that accommodated extension vnthin the Great Valley forearc concurrent vdth Tithonian-Valanginian through Coniandan sedimentation (ca. 151 - 86 Ma). Down plunge projection of the Great Valley monocline and interpretation of seismic reflection and borehole data reveals 228

that these faults collectively form a graben in which there is attenuation of pre-extensional rocks (Coast Range ophiolite and Stony Creek Formation) and

stratigraphic expansion of S5niextensional imits in the hanging WciU. Footwall rocks show the opposite relationships, pre-extensional strata are thick and synextensional strata are thin compared to the hanging wall. Tectonic transport of the hanging wall was to the northwest, hence entire biostratigraphic zones from both the pre-extensional and synextensional rock packages are missing because they were displaced and later eroded coeval with development of the Great Valley monocline.

2) West-dipping, shingled reflectivity beneath the west and central Great Valley Group is interpreted as interlayered mafic volcanic flows, volcaniclastic rocks and intrusives of the Coast Range ophiolite. This conclusion is based on correlation of ophiolitic rock in the subsurface to exposures of the Coast Range ophiolite and seismic reflection analogs from the Atlantic margin from this study combined with others investigation of borehole samples, gravity, magnetic, and seismic refraction data. Complete repetition of the lower Great Valley by a bUnd Franciscan-wedge sole fault commensurate with 40-50 km of horizonteil shortening is precluded (e.g., Wakabayashi and Uimih, 1995).

3) Dip divergences related to the subsurface feature we have termed a "fork structure" reflect a combination of sjmdepositional normal faulting and stratal onlap along the flank of the evolving Great Valley foreeirc basin, rather than post-depositional thrust wedging. Late-stage diastrophic mechanisms 229 are not required to interpret a forearc that owes much of its present-day bedding architecture to processes coeval with deposition. 230

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