Extensional tectonics of the Cordilleranforeland fold and thrust belt and the Jurassic-

Cretaceous Great Valley forearc basin

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Authors Constenius, Kurt Norman, 1957-

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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

UMI Microform 9829328 Copyright 1998, by UMI Company. All rights reserved.

This microform edition is protected against unauthorized copying under Title 17, United States Code.

UMI 300 North Zeeb Road Ann Arbor, MI 48103

2

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

Roy A. Jyohnsori

Peter J. Coney,' -s.

P Peter G. DeCelles

Susan L. Beck

Doctor of Philosophy

Date f i ^

1 Date

Date

•fA<//9<F Date

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

ector _:£Zj 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

1 0

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

1 2

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

1 5

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.

1 6

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

1 7

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

390N^

BRiTlSH COLUMBIA ALBERTA llftOH* iiyw lU^W ntw'

F CASADA Kishenehn

bastn Bearpaw mountains

Whitefisft ^

Kaiispell

Flathead Lake Htghwood

Mountains Great

Falls

Missoula Eastern Limit

Heltna

if " (j Buttt

^ afi'

Crazy Mountains

Boztman

Horse Prairie i5»Nf— & Medicine Lodge

basins

East-Central Idaho half grabens

Jackson

Idaho Falls IDAHO

• Poeatello

200 km 1

Fotvkes & Woodruff half grabens

Great

Provo • UTAH

4I<>Nf-

MONTANA

Thin-Skinned Thrusting

Southwest Montana Province of Basement Involved Thrusts

(1) Kishenehn basin 12) Tobacco Plains Valley (3) Flathead Valley (4) South Fork basin (Hunm Horse) (4b) South Fork basin (While River)

Mission Valley Swan, Avon-Nevada Creek & Blackfoot Valleys Lincoln Valley Douglas Creek Valley Flint Creek Valley Deer Lodge Valley Helena tc. Townscnd Valleys Smith River Valley Toston Valley Qarkston basin Madison-Gallatin basin Norris basin Three Forks basin North Boulder basin Jefferson River Valley Lower Ruby River Valley Beaverhead Valley Divide Trcnch & Melrose Valley French Gulch basin Big Hole Valley Salmon Valley Grasshopper Valley Horse Pnirie basin Medicine Lodge basin Muddy Creek basin Sage (.reek & Red Rock River basins Upper Ruby River Valley Centennial Valley Lemhi tc Birch Creek Valleys East-Central Idaho half grabens Fowkes half fpraben Woodruff half graben Morgan Valley - Huntsville graben East Canyon half graben Farmington-Wasalch Front basin

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 Sanpete Valley Sevier Valley

2 0

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

2 2

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.

2 3

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! (Mlitdle Eocene-Early Miocene)

Middle Eocene Igneous Rocks Synorogenic Sedimentary Rocks

(Cretaceous-Early Eocene)

Igneous Dikes & Sills

Detachment Surface with bath Thrust and Normal Movement

Precambrlan Crystalline Basement

to

2 5

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

ALBERTA CANADA

KISHENEHN^ ~ Ji BASIN '^-

U.S.A. MONTANA

TOBACCO] VALLEY Vr Cenex 4-liS ^ \ L»d«nbnij

Willow Creek Fm

OBrowning

EASTEKN BOUNDARY OF

FOLD A THRUST BELT

AULT Whitefish

FLATHEADMi&::i VALLEY

Kalispell Sample of Hoffman et al. (1976)

SOUTH FORK

•iXBASlH Chateau o nspvi Flathead

Lake

Augusta MISSION FAULT 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

2 7

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.

KISHENEHN BASIN LEWiS THRUST SHEET Kilometer!

SlnicturilJUilcl r, FUlhead Ramp tR*R)

Ihicknfus of If wis thrust shrrt at inception of crustal ntension

I ocrnf I - 7'K Am

WILLOW CREEK FM

KISHENEHN FM liathead duplet

structural relief J km I Ysh-Ysnl!;i£:l!Si^ TIIKilSl flatheail ramp

structural relief km p^iAI. DETArHMKNT - AUTOCHTHONOUS CAMBRIAN-CRETACEOUS

ROCKS

WATERTON DUPLEX

ARCHEAN CRYSTALLINE BASEMENT

FLATHEAD DUPLEX

Kllomeleri

Is) 00

2 9

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

3 0

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

3 2

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

3 3

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

3 4

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:

3 5

(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<mibrian detachment horizon that dips conformably with autochthonous

basement ~ 2-3° (35-52 m/km) southwest (Yoos et al., 1991; Bally and others ,

1966). Displacement on the Flathead fault is - 13-15 km to the southwest

(Constenius, 1988; McMechan, 1981).

(3) the Kishenehn basin is a half- or asymmetric graben in the central

and northern parts of the basin. The dip of the Kishenehn Formation is ca.

40-500 in the lowest beds and decreases systematically upsection and toward

the Flathead fault to ca. 20-25° (Constenius, 1981; McMechan, 1981). Hence,

preserved in diese middle Eocene to early Miocene synextensional strata is a

record of -40-50° of total rotation of tfie hanging waU; -20-30° of growth-

fault-related rotation, and -20-25° post-depositional rotation. To the south in

the Middle Fork region (Fig. 2.5), strata of the Kishenehn Formation are

middle Eocene in age and have been rotated as much as 50°, 18° of which is

growth-fault-related rotation, and about 30° represents post-depositional

rotation.

3 6

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.

Cr/sial Creek

Middle Flathead River

to

tft a-5 a a* a to a c«

_ Wolftail A' Mountain

§ ^

3 8

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

3 9

(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

4 0

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.

4 1

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 Dry Ck

Lightning Ck Km Hope slock Moule / Ravalli-Missoula Area-Marathon

fault / thrust^/ Cpa " Gibhs

Purcell anticlinorium

FPRICWFI

Pinkham thrust

Jocky Kishettehtt Mountain

Trench South Fork basin Flathtad

Hefty thrust /"""

Lewis thrust sheet

Rl'

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

• • :

4^ N>

4 3

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

4 4

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

4 5

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

R2I

Idaho 1130W 1120W mow Qc«r Lake

Woodruff half grabea

WDF-1

Fawkes

Marfan VafUy HuntsvUic

Amoco State Uiah'L Bridge 1130W

410N-i-Wyomine iitpw —i " i 41 Porcupine

Ridge

East Caityan half trabett

Great Salt Lake basins Salt Uk Qty Farmington' Wasatefi Front basia

Traoerse Rante Keetley & Park C/ly

ooleanie fields volcanic fiel

Vbble half grabea Cettaawcod arch €r

iVnsatch intrustoe belt Sale

Ptovo Utth LaJce

Meadows Paysoa Lakes half grabens

Utile Diamond

Creek half grnben

mow

Charleston'Nebo thrust

little Clear Crttk half graben

(Volanlc it *0 IntnitWe - bljek)

rTYcambtian rocks Tintic Moaatatas volcanic field

X Thrust fault f Nornul rjuit Sanpete Valletf

50 km

^I^Fowkw Foniutioo fc Norwood Tuff

ONorwood Toff-Moroni FormaUon b cQUtvslcnlt

Ute eocene01lgo«ne Inieoof rock« ilcanldMtie • ttippic.

fuab Valley

4 7

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).

4 8

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

4 9

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

5 0

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.

Northwest Meters

3000

Fowkes half graben - Medicine Butte thrust Anschutz Ranch

Field Anschutz Ranch East Field ANSCHUTZ JJ-I

A N S C H U T Z 3-1

ANSCHUTZ A.R.E. ^.K.£.

VVII-OI , THOUSAND PEAKS E't

Southeast 28 ta2if-2 Meters

4 JOOO Acocks Nonnal Fault

Tw-Tev-KTei

Medicine Butte^ Thrust

Upper Cretaceous Rocks

Upper Cretaceous Rocks

1000

• Km

- 'tm

. -3000

« '400Q

•5000

to

5 3

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

5 4

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

5 5

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

5 6

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).

Northwfut

0 0

Vi

o u

GSL-3 Great Salt Lake basin

Amoco Utah-L A

VVrsI

Southeast ®®|~

WDF-l

Woodruff halfgrahen iMit

Neogene rocks -40j^r/^9Ar 10-11 IMa

10.3 +/- 1.0 Ma

.^uat^rnary luvium

./ Fowkes

10.5 fcm » SJI km Norwood Tuff

^ \ e q v . ,

Wasatch

0.5 km

^•0 km Listric normal

40/{r/39/{r 38,5 +/. 0.2 Ma

Fm .

Q

10

Northwrat

00

«: o 01 Vj

R2

Listric normal fault

Wasatch Culmination

EWK-1 Fowkes half grnben

' W a s a t c h F m

Fowkes

0.5 km 2.0 km

S SI

a. 10 u Q ...

Great Salt Lake basin -A

Morgan Valley Wasatch fault Huntsville graben

Evanston Fm

•Acocks normal Woodruff fault

half graben

V. Reactivated sal detachment J H M H ^ crystalline basement

.. .. . Fowkes -^""buiH l-'f g">""

5 8

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-

5 9

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

6 0

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.).

6 2

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

CRETACEOUS

PALEOCENE

EOCENE

OLIGOCENE

MIOCENE

PLIOCENE

49 47 45 43 LATITUDE

OF TECTONISM

M L

Thrusting -foreland basin sedimentation

Sevier belt collapse Metamorphic core complexes

Basin & Range extension

o o 41 39

PROTEROZOIC CRUST

6 4

(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

6 5

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

!:

- |S £;uj oB 5S z.< eSc Bt-i O" X u)

EASTERN LIMIT OF THRUSTING BOULDER HIGHWOOD

MOUNTAINS CRAZY BEARPAW

MOUNTAINS MOUNTAINS

Northeast

BLACK HILLS

NI-WrORT FAULT (TIGER FMI "

z o ui CO

zZH ^ I/) s s X X tlJ m " H

TIME OF THRUSTING

(-7J.S7 Ma)

(HOFFMAN ET AU 1976^

YOUNGEST DISPLACEMENT UN LEWIS THRUST (-65-55

(COVEY ET AU 19941

YOUNGEST

SEDIMENTARY ROCKS

DERIVED FROM THRUST BELT

(WASATCH FM^ -57-54

— KISHENEHN

BASIN

FOWKES h WOODRUFF

HAIF GRABENS

If

EASr-CENTRAL IDAHO

HALF GRABENS

AGES OF UNDEFORMED

INTRUSIVE ROCKS IN THRUST BELT*

(HARLAN ET AU 1900)

HORSE TRAIRIE-MEOICINE LODGE BASINS

I J- X -i-

300 400 500 600 700 800

DISTANCE FROM TRENCH (km) 900 1000 1100

6 9

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),

7 0

'rv

oc:^cp^'°^ !«. uSA.

CORDILLERAN FORELAND FOLD & THRUST BELT

rv KtSHENEHN BASIN

(48-21 Ma) PRIEST RIVER

(SS.42 HR,,"^T

«3-<S MA> "

HGW

M 'R O"*' •^OO' %p

E-^ST-CrvTB V ' / '^'NTRal IDaho > Hiirr & O'C

^'ON-RAFT RIVED "{SCAR'ARA "3-37 Man

RB-EHCW# MAJ

« FOWKES-WOODRUFF HALF GRABENS

P* (49-27 Ma)

SIERRAN-WASATCH • RANGE MAGMATIC BELT V (37-21 Ma)\ \ >• *

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

OF SUBDUCTED SLAB

1 -\

"ORTH

500 km

OCORDILLERAN FORELAND FOLD AND THRUST BELT

FORELAND PROVINCE

LATE PALEOCENE MAGMATIC BELTS

o \LATE PALEOCENE HALF GRABENS AND ' GRABENS IN THE THRUST BELT

METAMORPHIC CORE COMPLEX -C

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.

7 2

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

1 3

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

7 4

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;

7 5

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

7 6

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

7 7

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.

7 8

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.

7 9

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

8 0

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).

8 2

Wasatch fault

Salt Lake City • p V ^

Cottonwoo<d

Cottonwood I metamorphic| core complex'

(Figure 7)

Tibbie half-graben

Provo Little Diamond Creek

half-graben CN

Santaquin Meadovrs &

Payson Lakes huf-grabens

Keetle / volcanic

V r o t e r o ^ » * ^ g

111 ocw enjtfaulf 4-4(f30'N

Triangle zone

©

Uinta basin-mountain boundary fault (blind thrust)

Nebo thrust imbricate

Charleston- Nebo thrust f

N Little Clear Creek

half-graben 20 km

8 3

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

8 4

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

8 5

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,

8 6

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

Wamtch Fork Witmich fhuit MounMiu "ilf-graben Clear Creek Hollow

msu.,„n J"'f'l^oten anticline Emery uplift

CN' Vast

D f f r Cm* U i T u a t Dairy Fork jhult Plateau

Sohio Indianola # 7 trm>/Kt,d ?.OtmNE)

'^"'flinuloflriannlezow IhttUt / Cn

Xnjo/i/i finita

Uinla BMB thrust

Ihvl Z I f - c

Pile

«•

•"-5 km

- - 1 0 k m

km

00 -vj

8 8

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).

8 9

25 r

25 km

East

UINTA BMB THRUST

Santaquin culmination (-82-40 Ma, 20 km) West

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

Uinta BMB sheet Nebo thrust sheet Emery uplift

9 0

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.

9 2

Little Diamond Fork half-graben Little Diamond

Pork fault

Little Clear Creek half-graben

West Thistle fault

Rollover anticline

Union 1-9 (projected S.OhnNE)

Oil Hollow anticline

Hunt Inditmola #24-1 (pnjected S.OhnNNE)

Emery uplift

Wasatch Plateau

Dairy Fork fimlt Sohh Indianala tl (projected 7.0 km NE)

East

^0

10

Little Diamond Fork half-graben

Little Clear Creek half-graben

Little Diamond Fork fimlt

West \ Vdsiu

Rollover anticline Thistle fimlt

Oil Hollow anticline

Hunt tndianola #24-1 (projected 8.0kmNNE)

Dairy Fork fimlt Sotop ImUanoUt *1 (pmjectcd 7.0hnNE)

T4:

Emery uplift Wasatch Plateau

Union 1-9 (projected 8.0hnNE

^-N sole fault

9 3

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

9 4

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

9 5

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

9 7

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),

9 8

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

9 9

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

1 0 0

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

1 0 1

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.

c West

Bowl-shaped half-grabens Little Diamond Creek Little Clear Creek

half-graben half-graben

c East

Little Diamond Fork fault

Thistle Lower Moroni Fw fault U^erMoroniFm

40.9+/-1.9 Ma

Buried Hogback

fSOOO

-2000

-WOO

- s.l.

- -1000

-2000

o N>

1 0 3

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

1 0 4

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

1 0 5

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

1 0 6

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

1 0 8

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-

1 1 0

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

1 1 1

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 a a a a a a a / 'Keetley volcanic''

field X Ontario stock

(36.0 +/-10) Aha stock 5.0 */-10)

Clayton Peak stock

sAfSO/ ^ . ^2^

/ / / ^ f V / / / / / / / , #>> vX Vn V-^'Vj^wXS. 6 n ,

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

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

/0ftm

Wasatch fault (Salt Lake segment)

Comer Creek tectonite

Deer Creek detachment fault

Silver Lake fault

Wasatch fault (Provo segment)

Box Elder Peak anticline / Charleston-Nebo thrust Tibbie half-graben

1 1 3

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

1 1 4

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

1 1 5

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 South crestal collapse graben

American Fork

Forest Gate fault

A'

North Elevation

rSOOO Little Cottonwood stock

- 4000 Box Elder Peak Silver Lake Ridge

"2000 N S N \ rn* S N N / • / TV » • / N N N N l I N N S

Stiver Lake mylonite / / , , <

Deer Creek detachment fault system -^reactivated Charleston-Nebo thrust i

--woo Alta thrust

ON

1 1 7

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 Southwest

Mount Timpamgos

Box Elder Peak Q' anticline

^ , Northeast American Fork

crestal collapse graben Deer Creek aetachment fault system r 5000

Stiver Lake detachment Ttbble

half-graben '4000

- 3000

i.Uvi'.kK -2000

fm - 1000

i^iP

-1000

1 1 9

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.

1 2 0

1 2 1

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.

1 2 2

1 2 3

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

1 2 4

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).

Wasatch fault

I i5f| tU

pCs-CD

Ppu(Pw)

pCs-CD

Ar/Ar

/ a = 50 Fw/PPoq

205 n= 113

n=6l8 Pw/PPoq

218 II = 337

249 n = 62

North

10 km

Charleston-Nebo thrust

Deer Creek detachment Tibbie half-graben

.5 +/- 0.2 Ma

.5 +/-1.4 Ma Ar/Ar 36.4 +/- 0.2 Ma

25" 45° '

Porcupine Gulch dikes

Silver Lake detachment fault

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

I l l

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

1 2 8

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

1 2 9

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

Silver Lake mylonite (SL)

Wasatch Fault

littlf Cottonwood / stoc'^ ^

i i -l^tachment fault i system

Dry Creek Canyon (DY) shear zones (86)

Equal Area Corner Creek

tectonite

Deer Creek Pass (DP) shear zones (55)

Equal Area

Mill Fork (MF) shear zones (72)

Equal Area

Mul Canyon Peak (MC) shear zones (29)

Equal Area

• "-rjfsi. ? ...

o

1 3 1

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

o3 = 249",0\

a3 = 250",5" a3 = 252^20"

a3 = 230",5"

Dry Creek Canyon (DY) (T-axis 254'© 7°)

Equal Area

Utuc Cottonwood^

'jVa^tes Fms)

N

10 km

Phyllonite in Fort Canyon (FC)

L-S(270°.(f)

a3 = 233"5"

o3 = 255",5"

S-C(226°.2f) ^s-C

8 km displacement

S-C(12(f,4<f) ,5

i^jtiA^^orizontal offset in Big Cottonwood Fm)

Mill Canyon Peak (MC) (T-axts 239 @ t)

iiiipliii! Deer Creek Pass (DP) (T-axis 263°® - J)

Equal Area

Mill Fork (MF) (T-axis 20® f)

Equal Area Equal Area

:n: m

-Hiu Ifjljii?: •iTi; rthj;;.

Utf:::::]!?:

N)

1 3 3

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;

1 3 4

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

1 3 5

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

1 3 6

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).

1 3 7

Uttte Cottonwood stock

Figure 16

Silver I )^362m

\ 2740 m \ 0

SL2. \J.. SUverLake \

"^SLS detachmentfauU j

Folded dike

Aua -r , ; / 36 thru j- 20

3<mm JL

3199 m

Porcupine / >:Gulch

\dikes

292 m

ii-^SuasLQi 111 37 '

Deer Creek normal fault Tibb le half-graben TibBle half-eraben *• / R ?•"" J' 1

49 3(r-

1 3 8

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.

South

STRUCTURAL RELATIONSHIPS EXPOSED IN WEST WALL OF SILVER LAKE CIRQUE

Dikes

Marble tectonite

Silver Lake c , , , , detachment fault mylonite

Uttle Cottonwood stock

North

f| . \ \ N \ * » / / / / / /

•e 20-

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

20-

1 5

Mylonite Whole Rock

-i 0.0 02

Fraction

-I r-0.4 0.6 39

Aplite Dike _ Whole Rock rt 15 •'

g; 0.0 0.2 0.4 0.6 0,8 1.0

< Fraction Ar Released

15 0.8 1.0

'Ar Released

K-Feldspar

Fraction ^®Ar Released

1 4 0

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.

1 4 1

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

1 4 2

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

1 4 3

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).

1 4 4

LITTLE COTTONWOOD STOCK 3;; Snowbird & Silver Lake Composite

3< CD m CM « CO o CM

30.5 ± 0.5 Ma 207V235

LITTLE COTTON"WOOD STOCK: Silver Lake

30.5 + 0.6 Ma

LITTLE COTTONWOOD STOCK Snowbird 31

30.5 + 0.8 Ma

ALTA STOCK Brighton

33.5 ± 1.0 Ma 207V235

ALTA STOCK Guardsman Pass 40.

35.0 ± 2.0 Ma (?) 207*/235

CLAYTON PEAK STOCK 3i

00 m CM « CO o CM

35.5 ± 1.5 Ma 207V235

ONTARIO STOCK

TO

36.0 + 2.0 Ma 207V235

1 4 5

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, 111 0 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, 111 o 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

1 4 6

145-175 4 127 61 4.55 602 470 510 2.4 36.5 ± 0.5 36.2 ±1.1

+1 0

CM

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

80-125 4 22 13 2.23 166 212 323 2.8

CO

+ 1

CO

CO

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 0.14 i 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.

1 4 7

TABLE 4.2. SUMMARY OF WASATCH INTRUSIVE BELT AGE DATA

Lim J. COTTONWOOD STOCK - NW FT APATITE 85 *!- 1.0 rai FT ZIRCON 245 0.6 #CR FT SPHENE 24.1 •/- 0.6 #CR K/Ar BIOTITE 25.1 0.8 #CR K/Ar BIOTITE 24.61 •/- 054 #J K/ArBIOrrTE 25.15 •/-0.24 « K/Ar BIOTITE 25.83 •/- 0.70 « •^OArf 'Ar BIOTITE 28.4 •/- 05 PLATEAU #F •"•Aif 'Ar HORNBLENDE 303-<-/-2.2 PLATEAU #F K/Ar HORNBLENDE 31.1 •/-0.9 #CR K/Ar HORNBLENDE 29.93 -•/- 1.66 #J K/Ar HORNBLENDE 30.72 •/- 058 #J RWSr 24 *!- 6 #C1

LITTT J. COTTONWOOD STOCK - S ' FT APATITE 6.4 •/- 2.0 lo 10.8 *!• 2.8 #K FT ZIRCON 9J •/- 0.8 to 20.4 •/- 1.6 /K K/Ar SERICITE 17.6 •/- 0.7 t9 *°Ai/ 'Ar K-FELDSPAR 7-17 SPECTRA (li8 •/-0.6 INT) #C '*°/W 'Ar K-FELDSPAR 12-40 SPECTRA #C "^OAI/ 'AT BIOTITE 243 •/- IJ PLATEAU #F

ALTA STOCK . F^inORANULAR FT APATITE 215 *!• 2.0 #AKC FT APATTtE 34.3 •/- 15 #M FT APATITE 35.6 •/- 3.6 #€4 FT ZIRCON 35.0*/-2.9 #AKC FT ZIRCON 33.7 •/- 0.9 #CR FT SPHENE 32.7 •/- 0.9 #CR K/Ar BIOTITE 32.1 •/- l.O #CR K/Ar BIOTTTE 32.6 *1- 1.0 fCR K/Ar BIOTITE 30.17 •/-054 tS "•"ai 'AT BIOTITE 31.97*/-0.23 INT #J "' At^ 'Ar HORNBLENDE 33.72 +/- 050 INT #J K/Ar HORNBLENDE 335 •/- 1.0 #CR K/Ar HORNBLENDE 30.9 •/-15 ICR U/Pb ZIRCON (4) 335 */- 1.0 #C

CLAYTONFEAKSTOCK FT APATITE 16.9 +/- 2.0 #AKC FT ZIRCON 35.9 •/-3.1 #/VKC FT APATITE 41.2 •/-3.0 ICR FT ZIRCON 40.9 •/- 2.2 »C8

K-FELDSPAR 31-40 SPECTRA #C K/Ar BIOTITE 32.9 •/- 1.2 #B K/Ar BIOTITE 34.8 •/- 1J #B K/Ar BIOTTTE 36.7 */• 15 #CR ^°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 APATITE 35.1 •/- 3.0 #M K/Ar BIOTTTE 35.2 •/- 13 #B K/Ar BIOTTTE 36.8*/-1.1 ICR K/Ar BIOTTTE 35.63 *1- 0.32 « K/Ar HORNBLENDE 41.08 */- 052 « "•"Arf 'Ar HORNBLENDE 385 V-0.7 PLATEAU;

ISOCHRON 38.0 •/- 0.9 #F

VALEO STOCK K/Ar BIOTTTE 34.6 •/- 1.6 #B K/Ar HORNBLENDE 403 *1- 1.6 #B6 K/Ar HORNBLENDE 39.8 •/- 1.2 #B7

MAYFI^WKR STOCK FT APATTTE 38.9 •/-55 #M FT APATTTE 38.6 •/- 3.6 #/UCC FT ZIRCON 3Z5 •/-il #AKC K/Ar HORNBLENDE 41.2 •/- 1.6 #B ^°Ai/ '/Vr HORNBLENDE 32.7 •/- 1.4 PIATEAU IF

t rm F. COTTONWOOD STOCK - NF. FT APATTTE 10.2 »/- 1.2 lAKC FT ZIRCON 29.7 •/- 1.9 lAKC K/Ar BIOTTTE 273 •/- 0.9 ICR K/Ar BIOTTTE 27.99 -/- 036 IJ K/Ar BIOTTTE 29.10 •/- 3.40 IJ *°Ai/ 'Ar BIOTTTE 31.8+/-23 PLATEAU IF K/Ar HORNBLENDE 3153 •/-4.40 IJ U/Pb ZIRCON (4) 305 •/- 0.6 IC

LITTLE CQTTQNWOQP STOCK • ^VHITE ECiE K/Ar MUSCOVTTE 255 •/- 0.8 ICR K/Ar SERICTTE 23.41 •/- 0.46 IJ

t lTTIJ-: COTTON^VOOD STOCK - SF. FT APATTTE 4.9 •/- 0.8. 7.9 »/- 1.2 & 123 »/- 15 lAKC FT ZIRCON 29.8 •/- 1.2 lAKC ''Ai/ 'Ar WHOLE RX 27.0+/-03 PLATEAU [DIKE] IF

WHOLE RX 27.1 W-O.l PLATEAU[MYLONTTEJIF K-FELDSPAR 12-27 SPECTRA (2i8 •/-13 INT) 0C

*°Ai '/Vr K-FELDSPAR 20-40 SPECTRA IC •^^AJ 'AT K-FELDSPAR 28.6 •/-0.4 INTEGRATED IC ^°A«/"Ar BIOTTTE 283 •/- 15 PLATEAU IF U/Pb ZIRCON (8) 305 •/- 05 IC

ALTA STOCK - roRPHYRY

'*°Arf 'Ar K-FELDSPAR 34.4+/-03 INTEGRATED IC K/Ar BIOTTTE 31.7 *1- 1.0 IB K/Ar BIOTTTE 31.6*/- 1.0 IB K/Ar BIOTTTE 3il •/- 1.2 IB K/Ar BIOTTTE 314 *1- 1.0 ICR K/Ar BIOTTTE 31.17 V. 054 IJ K/Ar BIOTTTE 3X8 •/- 0.7 ICR K/Ar BIOTTTE 33.26 •/- 055 IJ •*°Ai/ 'Ar BIOTTTE 32.28 ••/-0.24 INT IJ K/Ar HORNBLENDE 35.1 *1- I.l ICR ^°Ai/ 'Ar HORNBLENDE 29.90 •/-0.41 INT IJ

ALTA STOCK - GUAKDSMAN FT APATTTE 175 •/- 1.9 lAKC FT ZIRCON 37.2 •/-3.0 lAKC ••"/W 'Ar K-FELDSPAR 34.4 •/-03 LNTEGRATED IC K/Ar BIOTTTE 31.7 •/- 1.0 IB K/Ar BIOTTTE 31.6+/- 1.0 IB K/Ar BIOTTTE 3il +/- 1.2 IB K/Ar BIOTTTE 31.17+/-054 IJ U/Pb ZIRCON (2) 35+/-2.0 IC

FT /^ATTTE 38.9 +/- 2.9 lAKC FT ZIRCON 35.7 +/- 2.0 lAKC K/Ar PLAGIOCLASE 35.6+/- 13 (DIKE) IB K/Ar PLAGIOOASE 35.7 *1- 13 (DIKE) IB K/Ar BIOTTTE 33.4 +/- 13 IB K/Ar BIOTTTE 34.0 +/- 1.1 IB K/Ar BIOTTTE 33.7 +/- 13 IB K/Ar BIOTTTE 34.5 +/- 1.4 IB K/Ar BIOTTTE 333 +/- 13 IB ''Aj/ 'Ar HORNBLENDE 36.6 +/- Z1 PIATEAU IF

U/Pb ZIRCON (8) 36.0+/-2.0 IC

FLAGSTAFF STOCK K/Ar HORNBLENDE 39.7 +/- 1.2 IB K/Ar HORNBLENDE 37.8 +/- 15 (Dike) IB

HORNBLENDE 40.8 +/-1.9 PLATEAU IF

PARK PREMIER STOCK Klfa BIomTE 33.9 +/- 1.2 IB K/Ar BIOTTTE 31.21 +/-0.80 IJ K/Ar HORNBLENDE 35.2 +/- 1.0 IB K/Ar HORNBLENDE 3Z26 +/- 0.27 IJ

1 4 8

INDIAN HOI.IjnW Pt.lTf;

K/Ar HORNBLENDE 36J »/- 1.3 #B K/Ar HORNBLENDE 36.1 •/- 1.3 #B

KKF.TT.KY VOI/-ANlr.S

K/ArBIOnTE 35.1 •/- 1.1 #CR K/ArBIOTlTE 32.7 •/- 1.0 #CR K/ArBIOnTE 34.0 *!- 1.0 ICR K/Ar BIOTTTE 33.9 •/- IJ #B K/Ar HORNBLENDE 36.4 •/- 1J #B ^"Arf 'Ar HORNBLENDE 38.S»/-11 PLATEAU #F VERTEBRATE FOSSILS -34-37 #N

IAMPROITF Sn J .S fMOON CYN • FRANnS. im ""at 'AT PHLOGOPTTE 35.21 •/-0.10 PUVTEAUfMB "•"aj 'AT PHLOGOPTTE 34.93 •/-0.12 PLATEAU #MB "^^Ai/^'at PHLOGOPrrE 34J8 •/-0.11 PLATEAU #MB •^^Ai/ 'Ar PHLOGOPTTE34.10 •/-035 PLATEAU»MB

CI JXCOE STOCK FT APATTTE 30.3 »/- 1.0 #M

TRAVERSE VQLCAiMCS K/Ar BIOTTTE 38J •/- l.l ICR

TIBBLE FORMATION •^"Arf^^Ar BIOTTTE 36.4 •/-0.2 PLATEAU (LOWER TIBBLE] #F "•"Arf 'Ar BIOTTTE 28.5 •/-0.2 PLATEAU [LTPERTIBBLEJ #F

WHITES CREEK DIKE (E. CRANDAU. Cm "•"ai/ 'AT PHLOGOPTTE 13.93 *!- 0.09 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.

1 4 9

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.

1 5 0

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

1 5 1

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

U/reZfiton

o 20

Fraction Ar Released

SILVER LAKE CIRQUE - PALEODEPTH ~ 6 km K-Feldspar (KNC62695-2)

SIl Vr.R l AKi: MYI.ONIn; - PAI EODI H ll ~ b km 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)

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

1 5 3

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

1 5 4

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

1 5 5

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.

1 5 6

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

1 5 7

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

1 5 8

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 South Calculated Bouguer Gravity

X' North

5 -180

-200 OL ^ gs -180

O P -240 o CO

Observed Bouguer Gravity

.• ............^ ,, ,

Uinta allochthon

Reference Column

Charleston-Neho allochthon

Uinta Mountains uplift

Hogsback & Absaroka thrust sheets

^•2500 I I,.. i-p'2520 V-266Q

p-2540 UfptrCmt

p-2670

M4mOwt 0-2870 2800 / \ \ s

province Wyomms province (Archean crust)

UanoffCrmi p-2970 v Vs\Vs''v ('Pr0ter020/c crust)

' P r . » \ \ \ \ s \ \ s \ \ \ s s s \ s \ s s \ \ s \ \ \ \ \ / / / / • / / / / / / / / / \ M . 1 . . / • / / / / / / / / • / / • / / • Mono N \

Mantle P'3300 Mantle p - 3300

Crustal ivelt

100 DISTANCE (km)

200

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).

1 6 0

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

1 6 1

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

1 6 2

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

1 6 3

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 CC (MIDDLE EOCENE) rC

South ,T- North Utnta Mountains uplift

Charleston-Nebo allochthon / Absaroka thrust

t

^ h ^ i'J 5 ;• "i i •'> -20"

Mono

/ PROTEROZOIC

CRUST

CRUSTAL WELT

ARCHEAN CRUST

~30 km N-S crustal shortening

COTTONWOOD CORE COMPLEX CC (PRESENT) CC

North South Deer Creek detachment

fault

Little Cottonwood stock

mm. EWilfffliifll

Mono

~11 km exhumation related to

flexural-isostatic compensation

dikes & sills derived from crustal melts

mafic dikes & sills in loiver crust

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 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 \ S S N \ S V \ \ >

EUMW^

Extensional model - Harms, et al., 1992

Thin-skinned thrust model - Suppe, 1978

EAST

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 RANGE FAULT

STUDY AREA

Redding

Red Bluff

u^Chico Paskenta

\B ^ \ V

Sutter Buttes

V \ \ < GREAT VALLEY

GROUP Sacramento

San Francisco

Scale 100 km

COAST RANGE OPHIOLTTE

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

1

40° N

3SP30'N

122 W

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

Red Bluff

/ AHA

Paskenta

MH i

J 51 Sites anticline

Figure 10

10 km

40° N

3Sf30'N

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

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

Stony Creek Depth \ * • -o, ^ Conversion Velocity

(4.2 km/s) Si ^ **-dL.

ca Gulf Merle Hamm #1

Humble Michael #2^

• 0 Shell Vilche §2-5 J • Humble Fredericksen §1

*o American Hunter Alvarez #1 I . I

\

A i X

JU Jii

2 3 4 VELOCITY (km/s)

0

2 ^

4 Q

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

i l l l l l l loreno Sh Mokelumne Fm

PimnSi Maastrichtian

Starky & Wiaters Ss Sacramento Sh

Kione Fm Campanian Forbes Fm

Dobbins Sh illiams CynmT Santonian

GiundaSs Funks Sh Sites Ss Yolo Sh Coniancian

Ttironian riskeCkSh

SaltCk Cgl Cenomanian

Lodoga Fm Albian

I I I DISCO

Aptian NTINU.

Bairemian

Hauterivian

Valangiman

Bemasian

Tithoman Kimmeridgian

CRO volcanopelagic Oxfordian

Callovian Coast Range Ophiolite Bathoman

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

ARICA BIGHT FOREARC, CHILE Arica basin

Detachment

WEST LUZON TROUGH, PHILLIPINES

Detachment

ALEUTIAN FOREARC, ALASKA

Atkin basin

Onlap

ANCIENT GREAT VALLEY FOREARC, CALIFORNIA Lower Crctaceons NW Sacramento basin

Detachment

4 s 3 km

SCALE

VE - 6:1

10 km

Arcward

SOUTH SHETLAND FOREARC, ANTARCTICA llesperides basin - A

Detachment (?)

SOUTH SHETLAND FOREARC, ANTARCTICA Hesperidrs basin - B

SOUTH SHETLAND FOREARC, ANTARCTICA

Trench Hesperides forearc basin - A s.l.

Accretionary prism

Uceanic crust

Trench fill

S-0 s ^ ^

SO km

SUNDA FOREARC, CENTRAL JAVA, INDONESIA

Trench Accretionary prism

ceanic crust

50 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.

HF Campanian JL

HM ... .. V Sites anttclme

Polee to Faults Fold Axes Paskenta fault zone

Elder Creek fault zone Santoniun -6- 6* Fruto synclin

iurvman

Cold Fork fault zone

Sulphur Springs fault

GMH

Cretaceous (7

rri wwi

Aptiaii Discontinuity

1^3 Blttterivinn-B^netnian

Vahngitti(itt-$eTTia^ian

TithoHian-KimmeridgiaH CpMSt Range opltiplite

Coast Range Fault

Quntemary-late Tertiary SofUpninn (Wifliams Cyn)

Cpniancian

Turonian

Cenomanian

Franciscan complex

JO km ^Klamath terrane

N) o vO

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 Seaward-dipping reflections in oceanic basement

^ 3 S

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".

«) -30

> -40 6 -50

^ -60

A wsw

(J -70

-80

-90 o

§ -100 Coast Range

Fault

Calculated Bouguer Gravity

Observed Bouguer Gravity

Sacramento Valley Seismic profile TX-5

GMH Great Valley forearc basin p-2600 77

(A = - 500) //

'///A . P-28S0

'{A = +200;

p-2250 (A"-400)

p • 2470 (A = - mO) /I - 2770

(A = +120)

p - 2S50 (A = -100)

Coast Range ophwlite p - 2950 (A = +300)

p - 2570 (A = -HO)

CRO mantle - 3250 (A = +350) Franciscan comvlex

KsW Sierran arc ^^\'p.2820<&^-M)

27u5 (A= -195}

Moho / ^ / ^ \ \ \ \ \ \ \ \ \ \ p - 2710 X' P • 2820 fA = - 480)

(A = -590)^?\^^^ / / / f / / /

s s s / s \ s \ \

Mantle 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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