Neogene Stratigraphy of the Fish Creek-Vallecito Section, Southern California : Implications for Early History of the Northern Gulf of California and Colorado Delta

Neogene stratigraphy of the Fish Creek-Vallecito section, southern California : implications for early history of the northern Gulf of California and Colorado Delta

Item Type Dissertation-Reproduction (electronic); text

Authors Winker, Charles David, 1952-

Publisher The University of Arizona.

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Link to Item http://hdl.handle.net/10150/191123 NEOGENE STRATIGRAPHY OF THE FISH CREEK-VALLECITO SECTION, SOUTHERN CALIFORNIA: IMPLICATIONS FOR EARLY HISTORY OF THE

NORTHERN GULF OF CALIFORNIA AND COLORADO DELTA

Charles David Winker

A Dissertation Submitted to the Faculty of the DEPARTMENT OF GEOSCIENCES

In Partial Fulfillment of the Requirements For the Degree of

DOCTOR OF PHILOSOPHY

In the Graduate College THE UNIVERSITY OF ARIZONA

1987 THE UNIVERSITY OF ARIZONA GRADUATE COLLEGE

As members of the Final Examination Committee, we certify that we have read the dissertation prepared by Charles David Winker Neogene stratigraphy entitled of the Fish Creek-Vallecito section, southern California: Implications for early history of the northern Gulf of California and Colorado Delta

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

ezzL/ 1z' ),e,e.-(ler Date I9S(o Date (j Ç1c (q6 Date Ll -1--/ Date

I2( C //YCK 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.

Date 49ril /0/ /.787

Date 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 bor- rowers 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 the material is in the interests of scholarship. In all other instances, however, permission must be obtained from the author.

SIGNED: ACKNOWLEDGMENTS

This study would not have been possible without the unwavering support and encouragement of co-director Susan M. Kidwell, who intro-

duced me to the study area, arranged major funding, freely shared

results of her own research in progress, and repeatedly recharged my served as enthusiasm for the project. Co-director William R. Dickinson

graduate adviser, arranged initial funding, provided work space and stages of the facilities, and contributed invaluable input at critical

study. Committee members Kidwell, Dickinson, Peter J. Coney, Robert during F. Butler, and Victor R. Baker were all a source of inspiration

my doctoral program. National Science Founda- Major funding for this study came from came from tion Grants EAR-8407740 and EAR 8606254. Additional funding Laboratory of Geotectonics, the Geological Society of America and the

Department of Geological Sciences, University of Arizona. at Anza-Borrego Desert State Park is Cooperation of personnel G. Hendrix gratefully acknowledged. District Superintendant James Park rangers Paul Remeika, Jim Meier, arranged for collecting permits. advice, and companionship in the and Tim Davis provided assistance, and observations of his field. Paul Remeika generously shared results (many of which are unpublished geologic investigations in Anza-Borrego accompanied me in the field, incorporated in this dissertation) and

iii iv including two trips to the Carrizo Impact Area where he led me to important localities.

Employees of the U. S. Gypsum quarry at Split Mountain provided

ready access to company property for field work, and gave a tour of the quarry. Project Engineer Greg Cravener provided data on the gypsum deposit.

Sean Hagerty of the Bureau of Land Management in El Centro

furnished information on aerial photo coverage and availability.

Wes Peirce and Tom McGarvin of the Arizona Bureau of Geology

arranged for me to examine cuttings of the Lower Colorado River Project (LCRP) wells.

Ruperto Laniz of Palo Alto, California, prepared excellent thin

sections.

Marc B. Edwards visited the field area and contributed many

useful suggestions and observations on sedimentology.

J. Michael Boyles provided free copies of his nifty section-

logging forms.

Alison Hess reduced some of the paleocurrent data.

Richard S. Brokaw of Serious Design, Tucson, did some of the

drafting, provided graphic advice and burrito consulting, and generally

helped keep things at an appropriately serious level.

Steve at Willys Works and Dave at Luna Rebuilders provided

timely advice without which much of my field seasons would have been

spent immobilized.

Fellow Imperialists Nancy Beckvar, Carrie Campen-Bell, Ray

Bell, Alison Hess, Jim Lombard, and Dick Norris, Susan Kidwell's Advanced Stratigraphy class, and Jim Winker (my dad) shared many enjoy- able and occasionally productive hours in the Great Stinking Desert.

The Big A Gang (Rich Armin, Steve May, and sometimes Steve Wust,

Margaret Klute, and Chen Quanmao) inspired uninhibited tectonic theorizing.

In addition to those previously mentioned, the following people contributed reprints, preprints, theses, unpublished data, stimulating discussions, and/or correspondence: Mervin J. Bartholomew, Mark C.

Boehm, Anna Buising, Gary Calderone, John C. Crowell, Paul E. Damon,

Marlene Dean, Wilfred A. Elders, R. Gordon Gastil, Donn S. Gorsline,

James C. Ingle, Wallace A. Jensky, Noye M. Johnson, Dennis R. Kerr,

Larry Lemke, Everett H. Lindsay, Jon C. Matti, George Miller, John

A. Minch, Kevin Moodie, Charles L. Powell II, Judith Terry Smith,

Finn `lurlyk, Peter Van de Kamp, and Joan I. Winterer.

I wish to thank Jerry Lafayette and Beth Reilly, Jim Lombard

and Kik Moore, and Scott Brockenbrough and Jenny Shook for Wild Turkey

and refuge from the daily grind.

Finally, I must acknowledge my mentor James D. Howard, not only demonstrating that for getting me into this business, but also for

there is in fact life after geology. TABLE OF CONTENTS

LIST OF ILLUSTRATIONS

LIST OF TABLES

ABSTRACT

1. INTRODUCTION 1 Regional setting 10 Salton Trough 10 Gulf of California 14 Colorado River and delta 17 Climate 23 Basement geology 24 Previous work on Neogene stratigraphy of the western Salton Trough 26 Study area 37 Structure 39 Exposures 40 Logistics 40 Methods 42 Mapping 43 Measured sections 43 Paleocurrent measurement 44 Thin-section petrography 44 Geographic nomenclature and numbering scheme 46

2. SANDSTONE COMPOSITION AND PROVENANCE 48

Field observations 48 Thin-section petrography 50 Framework-grain composition 51 Other characteristics 56 Diagenesis 57 Comparison with other studies 59 Provenance suites 67

v i vii TABLE OF CONTENTS--Continued

Page 3. STRATIGRAPHY OF THE LOWER NONMARINE UNITS 74 Red Rock member (Mr) (new name) Sedimentary facies 75 Provenance and paleocurrents 79 81 Lithologically equivalent units Alverson 84 Formation (Ma) 84 Lithofacies Geochronology 90 and geochemistry 91 Lithologically equivalent units 92 Elephant Trees member (Me) (new name) Lower 93 fanglomerate unit (Mel), Split Mountain 94 Upper fanglomerate-sandstone unit (Meu), Split Mountain 98 Other fanglomerates 99 Lower Boulder Bed (lb) 100 Sequence in Carrizo Impact Area 102 Lithologically equivalent units 103 Fish Creek Gypsum (Mf) 104

4. STRATIGRAPHY OF THE MARINE UNITS 106

Latrania member (MPla) 108 Paleontology 115 Lithologically equivalent units 116 Lycium member (MP1) (reassigned name) 118 Transitional unit (MP1t) 119 Lower (thin-bedded) turbidite unit (MP11) 121 Upper turbidite unit (MPlu) 122 Paleontology and biostratigraphy 130 Kamia Creek member (MPk) (new name) 132 Stone Wash member (MPs) (new name) 134 Proximal facies 136 Distal facies 138 Upper Boulder Bed (ub) 138 Wind Caves member (Pw) (new name) 140 Sedimentary facies 148 Paleocurrents 150 Paleontology 152 Lithologically equivalent units 153 Mud Hills member (Pm) (new name) 155 Sedimentary facies 157 Paleontology 158 Lithologically equivalent units 159 viii TABLE OF CONTENTS--Continued

Page Deguynos member (Pdg) 160 Sedimentary facies and paleocurrents 169 Paleontology 174 Lithologically equivalent units 177 Lavender Canyon member (Plc) (new name) 178 Camels Head member (Pch) 181 Sedimentary facies 188 Paleontology 190 Lithologically equivalent units 192 Jackson Fork member (Pj) (new name) 193 Sedimentary facies 193 Paleontology 198

5. STRATIGRAPHY OF THE UPPER NONMARINE UNITS 200

Diablo member (Pd) 201 Sedimentary facies 204 Paleocurrents 216 Paleontology 218 Lithologically equivalent units 219 011a member (Po) (new name) 221 Sedimentary facies and paleocurrents 227 Paleontology 229 Lithologically equivalent units 231 Canebrake Conglomerate (PQc) 232 Lithologically equivalent units 234 Hueso member (Pa) 236 Sedimentary facies 237 Paleontology 238 Lithologically equivalent units 239 Tapiado member (Pt) 239 Lithologically equivalent units 241 Borrego (Pb) and Brawley (Qb) Formations 241 Paleontology 242

6. INTERPRETATION OF FISH CREEK-VALLECITO SECTION 244

Overview 244 Fence diagram 248 Proximality and paleotopography 249 Age control 251 Pre-rift phase 255 ix TABLE OF CONTENTS--Continued

Page Syn-rift nonmarine phase 259 Volcanism 259 Fanglomerates 261 Lower Boulder Bed 262 Unconformities 263 Pre-marine evaporites 264 Pre-deltaic marine phase 265 Fossiliferous shallow-marine facies 265 "Turbiditoid" facies 268 Upper Boulder Bed 278 Wind Caves submarine fan 279 Summary of pre-deltaic marine phase 281 Deltaic sequence 282 Prodelta 287 Delta-front cycles 288 Transitional (tidal flat) sequence 295 Delta plain 299 Paleoecological problems in the deltaic sequence 304 Summary of the deltaic sequence 305 Lacustrine units 308 Basin-margin alluvium 311 Margin of delta plain 311 Basin-margin fanglomerates 317 Basin-margin tidal flats 317 Summary 318

7. REGIONAL STRATIGRAPHIC AND CHRONOLOGIC SYNTHESIS 322

Genetic-stratigraphic units of Salton Trough 331 Unit 1. Pre-rift nonmarine clastics 331 Unit 2. Syn-rift nonmarine rocks 331 Unit 3. Pre-deltaic marine units 332 Unit 4. Deltaic marine units 332 Unit 5. Delta plain 332 Unit 6. Lacustrine deposits 333 Unit 7. Basin-margin clastics 333 Unit 8. Colorado terrace deposits 333 TABLE OF CONTENTS--Continued

Page Local summaries 333 San Felipe Hills, Borrego Badlands, and S. Santa Rosa Mountains 333 Yuha Basin 334 Jacumba area and Volcanic Hills 334 Sierra de los Cucapas 335 Cerro Prieto Field 336 Superstition Mountain 337 Whitewater Canyon and vicinity 337 Indio Hills 338 Durmid Hill and vicinity 338 Fortuna and San Luis Basins 339 Bouse Embayment 339 Possible extensions of Bouse Embayment 340 Circum-Gulf localities 343 Sierra San Felipe and Santa Rosa 344 Isla Tibur& 344 "Hermosillo Basin" (new name) 344 Coastal Baja California 345 San Jose del Cabo Trough 345 Cabo Corrientes 346 Tres Marias Islands 346 Mouth of Gulf 346 Regional chronology 347 Salton Trough and Bouse Embayment 347 Circum-Gulf of California 350

8. TECTONIC DISPLACEMENTS, PALINSPASTIC RECONSTRUCTION, AND PALEOGEOGRAPHY 354

Reference model 355 Constraints on tectonic displacements 357 Pacific-North American plate motion 357 Spreading at mouth of Gulf 359 San Andreas and San Gabriel faults 363 San Gorgonio Pass region 369 San Jacinto fault 370 Elsinore fault 373 Algodones fault 376 Salton Trough region 377 Agua Blanca fault 377 Displacements vs. rates 378 x i TABLE OF CONTENTS--Continued

Page Analysis of displacement histories Block 7 380 380 Block 6 386 Block 5 Block 386 4 387 Block 3 387 Effects of additional distributed shear 388 Trajectories 390 Paleogeographic maps 391 Palinspastic base 391 Faults and spreading centers 391 Marine transgression and regression 394 Tectonic-paleogeographic history 398 Plate-tectonic context 398 Proto-Gulf rifting 401 Transform plate boundary 402 Western Salton Trough 403 Pre-deltaic northern Gulf 404 Colorado River and delta 405 Lacustrine basin 407 Paleoclimates 408 Stratigraphic implications of this study 409

APPENDIX I: AERIAL PHOTOGRAPHIC COVERAGE 411

APPENDIX II: PALEOCURRENT DATA 418

APPENDIX III: SANDSTONE PETROGRAPHIC DATA 440

APPENDIX IV: DATA FOR PLATES 1-4 453

REFERENCES CITED 461 LIST OF ILLUSTRATIONS

Page

Figure

1.1. Index map to tectonic setting and Neogene localities of Salton Trough 2

1.2. Map of Neogene units in the Salton Trough and Bouse Embayment 3

1.3. Index map to geographic names in main study area 7

1.4. Index map to geographic names in lower Fish Creek drainage basin 8

1.5. Geologic map of Fish Creek-Vallecito (FCV) section 9

1.6. Schematic fence diagram of FCV section, Carrizo Impact Area, and Coyote Mountains 27

1.7. Stratigraphic units and symbols used in this report . 29

1.8. Stratigraphic synonymy for eastern Coyote Mountains and southwestern flank of Fish Creek Mountains (Carrizo 30 Impact Area)

1.9. Stratigraphic synonymy for FCV section 31

synonymy for Split Mountain west of Gorge 1.10. Stratigraphic 32 and southeastern flank of Vallecito Mountains

1.11. Stratigraphic synonymy for San Felipe Hills, Borrego southeastern flank of Santa Rosa Badlands, and 33 Mountains 38 1.12. Geologic map of lower Fish Creek drainage basin 45 1.13. Explanation for measured-section logs FCV section 52 2.1. Framework-grain composition of sandstones from of sandstones from 2.2. Miscellaneous compositional parameters 53 FCV section

x i i

LIST OF ILLUSTRATIONS--Continued

Page

Figure

2.3. Miscellaneous composition parameters for "C"-suite sandstones 54

2.4. Sandstone porosity and cement in FCV section 60

2.5. QFL compositional fields from this study compared with those of previous studies 62

2.6. QmFLt and QmPE compositional fields from this study compared with those of previous studies 64

2.7. Criteria for recognition of Colorado River-derived sand 68

2.8. Possible sources of discrepancies in sandstone compositions between this and previous studies 71

2.9. Geologic map of FCV section, showing distribution of provenance suites 72

2.10. Fence diagram of study area, showing distribution of provenance suites 73

and schematic cross section of Split Mountain 3.1. Geologic map 76 and southwestern flank of Fish Creek Mountains 82 3.2. Paleocurrent roses for Red Rock member Formation, 3.3. Index map of four outcrop areas of Alverson 86 in vicinity of Fossil Canyon, Coyote Mountains of the Alverson Formation and 3.4. Stratigraphic columns 87 bracketing units in part of Coyote Mountains

columns of the Alverson Formation and 3.5. Stratigraphic 88 bracketing units in Painted Gorge and vicinity

columns of the Alverson Formation and 3.6. Stratigraphic of Red Rock Canyon . 89 bracketing units in the vicinity Elephant Trees member, lower 3.7. Paleocurrent roses for 97 fanglomerate unit xiv LIST OF ILLUSTRATIONS--Continued

Page Figure

4.1. Details of stratigraphic relationships in and adjacent to Split Mountain Gorge 107

4.2. Index map of eastern and central Coyote Mountains 110

4.3. Stratigraphic columns of Latrania member 113

4.4. Graphic log of measured section SM-2, Split Mountain Gorge 120

4.5. Paleocurrent roses for "L"-suite marine units 123

4.6. Graphic log of measured section SM-3, Split Mountain Gorge . 124

4.7. Graphic log of measured section LW-1, Lycium Canyon . . 125

4.8. Graphic log of measured section 0S-2, Oyster Shell Canyon #3 126

4.9. Graphic log of measured section 0S-1, Oyster Shell Canyon #3 127

4.10. Graphic log of measured section KC-1, Kamia Creek, CIA . . 133

4.11. Stereonet plot of axes of outcrop-scale folds in Lycium member below Upper Boulder Bed 141

4.12. Graphic log of lower part of measured section NF-1, North Fork 142

4.13. Graphic log of upper part of measured section NF-1, North Fork 143

4.14. Graphic log of measured section SM-1, head of Split Mountain Gorge 144

4.15. Graphic log of measured section FC-5, Fish Creek Wash . . 145

4.16. Graphic log of lower part of measured section EK-1, 146 Cairn Canyon

Graphic log of upper part of measured section EK-1, 4.17. 147 Cairn Canyon XV LIST OF ILLUSTRATIONS--Continued

Page Figure

4.18. Paleocurrent roses for Wind Caves member 151 4.19. Graphic log of measured section FC-7, Fish Creek Wash . 162

4.20. Graphic log of measured section FC-1, Fish Creek Wash . 163

4.21. Graphic log of measured section FC-la, Fish Creek Wash 164

4.22. Graphic log of measured section NF-3, North Fork 165

4.23. Graphic log of measured sections NF-4 and NF-4a, North Fork 166

4.24. Graphic log of measured section NF-5, North Fork 167

4.25. Graphic log of measured section EK-2, Elephant Knees . . . 168

4.26. Paleocurrent roses for Deguynos, Lavender Canyon, and Camels Head members 173

4.27. Graphic log of measured section RR-1, Red Rock Canyon . . 180

4.28. Graphic log of measured section FC-2, Fish Creek Wash . 182

4.29. Graphic log of lower part of measured section FC-3, Fish Creek Wash 183

4.30. Graphic log of upper part of measured section FC-3 and section FC-3a, Fish Creek Wash 184 185 4.31. Graphic log of measured section NF-6, North Fork

of lower part of measured section NF-7, 4.32. Graphic log 186 North Fork

upper part ofmeasured section NF-7, 4.33. Graphic log of 187 North Fork

measured sections CT-1 and CT-la, 4.34. Graphic log of 194 Jackson Fork

lower part of measured section CT-2, 4.35. Graphic log of 195 Jackson Fork xv i LIST OF ILLUSTRATIONS--Continued

Page Figure

4.36. Graphic log of upper part of measured section CT-2, Jackson Fork 196 5.1. Simplified geologic map of FCV section, showing paleocurrent roses for "C"-suite sandstone bodies in Diablo and 011a members 202 5.2. Schematic cross sections of upper nonmarine units in western Salton Trough 203 5.3. Graphic log of measured section FC-4, Fish Creek Wash . 205

5.4. Graphic log of measured section FC-4a, Fish Creek Wash . 206

5.5. Graphic log of measured section FC-4b, Fish Creek Wash . 207

5.6. Graphic logs of measured sections DC-1 to DC-4, Diablo Canyon 208

5.7. Graphic logs of measured sections DC-5 to DC-7a, Diablo Canyon 209

5.8. Graphic log of lower part of measured section DC-7b, Diablo Canyon 210

5.9. Graphic log of upper part of measured section DC-7b, Diablo Canyon 211

5.10. Graphic log of measured section AT-1, Arroyo Tapiado . . 212

5.11. Graphic log of lower part of measured section LC-1, Layer Cake Wash 222

5.12. Graphic log of upper part of measured section LC-1, Layer Cake Wash 223

5.13. Graphic log of measured sections 0W-1 to 0W-5, 011a Wash 224

5.14. Graphic log of measured sections 0W-6 to 0W-8, 011a Wash 225

5.15. Paleocurrent rose for "L"-suite sandstones, 011a member . . 230 xvii LIST OF ILLUSTRATIONS--Continued

Page Figure

6.1. Summary of paleocurrent data for "C"-suite stratigraphic units in the FCV section 245 6.2. Genetic-stratigraphic units of the FCV section and

adjacent areas 246

6.3. Magnetic-polarity stratigraphy of Imperial and Palm Spring Formations in FCV section 253

6.4. Cumulative sedimentation and sedimentation rates for the FCV section, based on magnetostratigraphy 256

6.5. Summary of the pre-rift sequence 258

6.6. Summary of the syn-rift nonmarine sequence 260

6.7. Summary of the pre-deltaic marine sequence 266

6.8. Schematic representation of lateral change in a sandstone-conglomerate couplet relative to transport direction 270

6.9. Schematic representation and interpretation of retrogradational sigmoid geometry in upper turbidite unit of Lycium member 273

6.10. Summary of deltaic sequence, FCV section 284

6.11. Representative profiles from Meckel (1975) of modern Colorado delta, based on coring 285

6.12. Representative profiles from Van de Kamp (1973) of modern Colorado delta plain, Salton Sea flank, based on coring 286

6.13. Measured-section logs of delta-front cycles in upper Mud Hills and lower DeGuynos members 289

6.14. Profiles of Colorado River channel at Yuma gauging station, at maximum and minimum discharge 302

6.15. Distribution of Diablo and 011a members and Canebrake Conglomerate in western Salton Trough 314 LIST OF ILLUSTRATIONS--Continued

Page Figure

6.16. Suggested palinspstic reconstructions of Figure 6.15 . . . . 315

7.1. Neogene stratigraphie summary of Salton Trough based on genetic-stratigraphic units 323

7.2. Stratigraphie synonymy for Jacumba area 324

7.3. Stratigraphie synonymy for Sierra de los Cucapas 325

7.4. Stratigraphie synonymy for vicinity of Whitewater Canyon . . 326

7.5. Stratigraphic synonymy for Indio Hills 327

7.6. Stratigraphie synonymy for Durmid Hill 328

7.7. Stratigraphic synonymy for Fortuna and San Luis Basins . . 329

7.8. Stratigraphic synonymy for Bouse Embayment 330

7.9. Chronology of radiometrically dated rocks and events of circum-Gulf region 351

8.1. Neogene paleogeographic reconstruction of Salton Trough and northern Gulf of California based on palinspastic reference model 358

8.2. Age vs. distance for magnetic anomalies in mouth of Gulf of California 361

8.3. Schematic cross section of Ridge Basin, and sedimentary-tectonic chronology 366

8.4. Time-displacement diagrams for parts of Gulf of California-San Andreas transform system 368

8.5. Stereonet projection of virtual geomagnetic poles for FCV section 375

8.6. Total displacements vs. rates for components of the San Andreas fault system in the Salton Trough 379

8.7. Displacement history with uncertainty envelope for tectonic block 7 381 xix LIST OF ILLUSTRATIONS--Continued

Page Figure

8.8. Displacement histories with uncertainty envelopes for Agua Blanca fault and tectonic block 6 382

8.9. Displacement histories with uncertainty envelopes for Elsinore fault and tectonic block 5 383

8.10. Displacement histories with uncertainty envelopes for San Jacinto fault and tectonic block 4 384

8.11. Displacement history with uncertainty envelope for tectonic block 3 385

8.12. Palinspastic base maps for Miocene 392

8.13. Palinspastic base maps for Pliocene 393

8.14. Paleogeographic maps for Miocene 396

8.15. Paleogeographic maps for Pliocene 397

8.16. Timing of selected late Cenozoic events on the southwestern margin of North America 399

Plate 1. Tectonic map of the Pacific-North American plate boundary in pocket

Plate 2. Tectonic map of Salton Trough in pocket

Plate 3. Geologic map of western Salton Trough in pocket

Plate 4. Geologic map of Arroyo Tapiado, Carrizo Mtn. NE, and parts of Harper Canyon, Borrego Mtn. SE, and Sweeney Pass 7.5' Quadrangles in pocket

Plate 5. Measured sections of Lycium, Stone Wash, and Kamia Creek members in pocket

Plate 6. Measured sections of Wind Caves member in pocket

Plate 7. Measured sections of delta-front cycles in pocket

Plate 8. Measured sections of deltaic sequence in pocket LIST OF TABLES

Page Table

4.1. Tentative faunal list for Stone Wash member, Coral Wash exposures . . 137

4.2. Tentative faunal lists for Deguynos member 175

4.3. Tentative faunal lists for Camels Head member 191

4.4. Tentative faunal lists for Jackson Fork member 199

XX ABSTRACT

The Fish Creek-Vallecito section is the most stratigraphically complete and structurally intact Neogene exposure in the Salton Trough, and thus provides a useful reference section for regional stratigraphie revision and historical interpretation of the early Gulf of California

and Colorado Delta. The section comprises a marine sequence (Imperial

Formation) bracketed by nonmarine units (Split Mountain and Alverson

Formations below, Palm Spring Formation and Canebrake Conglomerate

above). Recognition of distinct suites of locally-derived and Colorado

River-derived sediment, combined with sedimentological evidence, led to

revision of this sequence in terms of informal members and genetic-

stratigraphic units: (1) pre-rift braided-stream deposits (2) syn-rift

fanglomerates and volcanics, with local pre-marine evaporites; (3) fan deltas; a pre-deltaic marine units, deposited primarily as small consisting of progradational sequence of the ancestral Colorado delta,

(4) an upward-shoaling marine sequence, and (5) a nonmarine delta- locally-derived basin- plain sequence; (6) lacustrine units; and (7)

margin alluvium that interfingers with (4), (5) and (6). mapping were Neogene palinspastic base maps for paleogeographic Pacific-North American plate based on displacement histories for the tectonic-sedimentary history boundary and its constituent faults. The Miocene rifting that propagated consists of: (1) early to middle mouth; (2) northward southward from southern California to the Gulf

xxi xxii marine transgression of the rift basin, reaching southern California by the late Miocene; (3) development of the San Andreas-Gulf of California transform boundary by inboard transfer of intraplate slip; (4) earliest

Pliocene initiation of the lower Colorado River and Delta by rapid epeirogenic uplift of the Bouse Embayment; and (5) late Pliocene or

Pleistocene transpressive uplift in the western Salton Trough caused by outboard transfer of slip from the San Andreas fault. The stratigraphic succession in the western Salton Trough resulted largely from tectonic transport through a series of paleoenvironments anchored to the North

American plate by the entry point of the Colorado River. CHAPTER 1

INTRODUCTION

Neogene stratigraphy of the Salton Trough region records a marine incursion from the early Gulf of California and subsequent initiation and progradation of the Colorado delta. These events were closely linked to extensional opening of the Gulf (Karig and Jensky

1972, Moore 1973, Moore and Curray 1982), initiation of the San An- dreas Gulf of California transform plate boundary and the Salton Trough

"leaky transform" hydrothermal system (Elders et al. 1972, Elders

1979), and epeirogenic uplift of the lower Colorado valley and southwestern Colorado Plateau (Lucchitta 1972, 1979). Thus, genesis and chronology of Neogene stratigraphic units in the Salton Trough have important implications for regional tectonic history.

The Colorado delta is the dominant depositional system in this sedimentary basin; it separates the nonmarine Salton Trough and the marine Gulf of California, which are otherwise a single, tectonically continuous transtensional basin (Figures 1.1, 1.2). The large influx of terrigenous sediment from the Colorado River has had a profound effect on the nature of crustal genesis at spreading centers of the Salton

Trough (Elders et al. 1972, Elders 1979, Muffler and Doe 1969). The modern delta has been the subject of a series of sedimentological studies (McKee 1939, Thompson 1968, 1976, Van de Kamp 1973, Meckel

116° 115° 114° edge of Salton Trough 35° key Neogene locality Needles active spreading center (pull - apart basin) major tectonic block

Whitewater Cyn.1(7, or Parker Indio Hills J\11)' 34° Riverside San "BA Gorge;nlo , 61/ San „erFelipe Hills & Blythe ",,,Borrego, Badlands 14,' A:cif- ,F.:é iscogi --:, e/yo 0 ,,,,,,,,,, ,Ç:O/ 1 4- 33° 5 CARGO ••• , • 's, MUCHACOH MTS ..--7San , ,,11— LAGUNA MIS Diego Y ulna

Yuha Basin

d') A.` ZON < oe• "••,80>, 32° .;> • 61.rp

MAJOR STRIKE-SLIP FAULTS California - San Andreas SA 31° MC - Mission Creek BA - Banning Sierra SJ - San Jacinto San Felipe EL - Elsinore CC - Coyote Creek AC - Agua Caliente SH - Superstition Hills SM - Superstition Mountain IM - Imperial LS - Laguna Salada AL - Algodones 30° CP - Cerro Prieto AB - Agua Blanca

50 100 150 km 9 50 100 mi

Figure 1.1. Index map to tectonic setting and Neogene localities of Salton Trough. Neogene of San Luis and Fortuna Basins is entirely in subsurface; at all other localities Neogene is exposed at surface. "Tectonic blocks" are discussed in Chapter 8. 3

1150

Pliocene lacustrine (Borrego Fm.)

Plio-Pleistocene nonmarine (Palm Spring Fm., Canebrake Cg., etc.)

Pliocene marine (Imperial Fm.)

Pliocene marginal marine (Bouse Fm.)

Bouse Fm. encountered In well

Miocene nonmarine (Split Mountain Fm., Alverson Fm., etc.)

major tectonic block

El Centro Yuma <

-c YUMA

+32°

MAJOR STRIKE-SLIP FAULTS

SA - San Andreas IM - Imperial AL - Algodones SH - Superstition Hills CC - Coyote Creek SM - Superstition Mountain AC - Ague Caliente nLS - Laguna Salada EL - Elsinore CP - Cerro Prieto Gulf of California

Figure 1.2. Map of Neogene units in the Salton Trough (exclusive of San Gorgonio Pass) and Bouse Rmbayment. Surface geology of Salton Trough is based largely on Dibblee (1954) and follows his nomenclature. Bouse Embayment localities after Olmsted et al. (1973), Metzger et al. (1973), and Metzger and Loeltz (1973). Colorado River distributaries after Merriam and Bandy (1965). 4

1976) and is one of the best-documented examples of a modern tide- dominated delta (classification of Galloway 1975). Colorado River sediment is compositionally distinctive (Van Andel 1964, Merriam and

Bandy 1965, Muffler and Doe 1968, Van de Kamp 1973) and has been

recognized in Neogene marine and nonmarine strata exposed along the western margin of the Trough (Merriam and Bandy 1965, Muffler and Doe 1968).

It might be expected, then, that the history of the Colorado

delta would figure prominently in stratigraphy of the Salton Trough.

However, most published stratigraphic subdivisions, nomenclature, and

descriptions have not explicitly acknowledged the presence, extent, or

significance of Colorado River-derived sediment in the Salton Trough.

Consequently, the ancestral Colorado delta cannot be separated from

other depositional systems on the basis of any existing framework of

formal stratigraphy. Major Objectives of this study, therefore, were to

recognize Colorado River sediment in the Neogene strata of the Salton

Trough (Chapter 2), to determine its extent in the stratigraphic sec-

tion and to define stratigraphic units that would reflect provenance

(Chapters 3-5), and to characterize the sedimentology of those units

(Chapter 6).

Marine faunas of the Neogene Imperial Formation in the Salton

Trough have been studied by paleontologists for more than a century

(Blake 1857, Arnold 1904, 1906, Vaughan 1904, 1906, 1917, Mendenhall 1978, 1910, Kew 1914, Hanna 1926, Woodring 1932, Bramkamp 1935, Foster

1979, Shremp 1981), especially because they provide evidence for 5 biogeography and paleoecology of the early Gulf of California (Durham

1950, 1954, Durham and Allison 1960, Stump 1972, Ingle 1974, Quinn and

Cronin 1984). The age of the Imperial Formation and, by implication, the timing of marine transgression have been a subject of controversy since the early studies; this problem has not been satisfactorily resolved. The Imperial faunas are also being studied now from the standpoint of taphonomy and paleoecology (S. M. Kidwell 1985, oral communication). Objectives of the present study in relation to paleontology and biostratigraphy were to clarify the physical stratigraphic relationships of fossiliferous marine units (Chapter 4), to provide sedimentological data to complement paleoecological evidence (Chapter

6), and to compile geochronological data for the Salton Trough and determine the regional significance of these data (Chapter 7).

Neogene paleogeographic reconstructions of the Salton Trough and Gulf of California are hindered by the complexities of syndepositional strike-slip faulting along the San Andreas-Gulf of California transform boundary between the North American and Pacific plates (Plate

1; Figures 1.1, 1.2). This boundary has accommodated at least 300 km of total dextral slip since the middle Miocene (Crowell 1975a, Moore and

Curray 1982). In the Salton Trough region, this total displacement has been accommodated by several major faults (Figures 1.1, 1.2; Plate 2), all of which were probably active during the late Cenozoic (Crowell and

Ramirez 1979). One problem addressed by this study was the construction on of palinspastic base maps which would honor available constraints 6 tectonic displacements while acknowledging the uncertainty inherent in such reconstructions (Chapter 8).

Stratigraphic and facies relationships in the Salton Trough are generally Obscured in the field by extensive structural deformation, which includes folding and faulting associated with wrench tectonics

(Plate 2). An exception to this prOblem is the Fish Creek-Vallecito

(FCV) section (Figures 1.3, 1.4) on the western margin of the Salton

Trough (Figure 1.1). The FCV section is essentially a coherent southwest-tilted fault block which provides an exceptional opportunity to

observe some 5-6 km of Miocene to Pleistocene basin fill exposed more

or less continuously within an area slightly larger than a 7.5' quad-

rangle (Figure 1.5). The FCV section, located within Anza-Borrego

Desert State Park, is the most structurally intact and stratigraphie-

ally complete exposed Neogene section in the Salton Trough region. It

contains a marine section bracketed by nonmarine units and it includes

the well known section at Split Mountain Gorge (Robinson and Threet

1974, Kerr, Pappajohn, and Peterson 1979). Magnetostratigraphy of part

of the FCV section by Johnson et al. (1983) provides the most precise

age control of any Neogene section in the Salton Trough region. In

addition, the FCV section is one of the areas in which Colorado River-

derived sediment had previously been identified in Neogene strata

(Merriam and Bandy 1965, Muffler and Doe 1968). The adjacent Carrizo

Impact Area (CIA) and Coyote Mountains contain historically important

fossiliferous exposures of the Imperial Formation. For these reasons,

the FCV locality provides a unique reference section and

115°05' 1113° 00' .... HARPER CANYON — I 7.5' OUAD TO BORREGO MTN. SE — 33* __L/I 7.6' OUAD 05' Harper • Flat .1 • ELEPHANT TREES VALLECITO 9... RAPIDER STATION 5, ( f ASK. . /0 ....,.„,---,----\.„-,\LNaeo Elephant I ..4 \_...... _:\..r7.es .. See. I Se Dave McCain Area • Spring (dry)

+ - - - 4. - -- ce-,L---\.:- .-- - - 6, . iiV + MOUNTAINS

-1'.1.? i; sk,e.to '4'46 Q. \ U.S. GYPSUM / QUARRY 4A.,.,ko, 7.7 4, I \I : \ °FISH CREEK.0 - , Off a • PRIMITIVE . r CAMP I (I ? ‘ 4o g : k I ) 1,t,... A' r : I 1.----1.-- area of Figure 1.4 g Z ir, ''-''k', fl. l. 1 . ---x,„.. . ' I 33. I \--,,,,...›'-1/4,1 —1_,,i. EL 00' 6) North'. ... Os, I"-, For* 4.5 •• eto,,,, 5 I -,-.--\. FISH -\''''\---,1• I ••.%4*b C"./.. *lash .../.4)&,/, m4 vot..19,---- ....7'7Nest` ,... ? Mud ,..) A 1.7 o• Hills F Ae, 4,e I '1- / CS), '' °'• ..m a* wae,:, 00y..4' ..1,e. 1/Boundary C R E Mountain ,,,Neg.6 East I I o - a /e /((./ Diablo Drop - off S' 1 -... I .. ri .8, view of / We:h. ''' Valley ° Badlands ••!.., 4 M T S. ). 710 _/,/) Mesa Middle m .0 / ° z Mesa .. 3 r..) ...•._ 01_ 1-\\siLt...s.„Vase.. .1 • I , Spring G 32° e Ir 55' v ? South - .., Palm . Spring Mesa .../ to riF. T urn-off 4 - so ---.....„ C„TIERRA M l 0 BLANCA

MTS.

CARRIE° MTN. NE ARROYO TAPIADO 7.5' OUAD 7.5' QUAD .40.- f -gar ------.... —41111W C eekier .4•"‘ Carrion I e Mountain I Palm • .5, Springs Carrizo Impact Area ----1 T ---' I ' .Z.4. c / ) .5.. co .. i Willov. A I\i .._ 4_ WASH CLOSED OR PAVED ROAD --- IMPASSABLE of soW wiL!ow ' • \,.., ,e :::y ENTRANCE TO rRANGER STATION El O. '"-.., SW,.. UNPAVED ROAD 32° , _ .--.• C 0 , _ ---^--... CANYON - --.•• 0 r e 50' ) i \ JACUMBA a Badlands 4, ye. 0 1 2 3 mi 2 . MTS. . • overlook 1 i I eli Ocot/1//ep 0 1 2 3 4 5 km c) s, lit] I I SWEENEY PASS • 7.6' QUAD ''''''''•••....eqfjen. t

Figure 1.3. Index map to geographic names in main study area (Plate in Figure 4). Names for eastern and central Coyote mountains are shown 4.2.

'00 le pedwi 03 oBarG un

1.114TAIN <,, c x a. ul ‹o of cc a#1.:7 0 0 0) 0 U.I w e+/— (/) _i 0 m 0Y -I < ---.------.., 4‘e' 0 co _ o.s. cyn. # ..:,--x J\5 ç S a) \ ../ GO g Ly 0002fris rj---..., 2 (:;.C..- - 4-0,.0- \ .S` o- ,..r,-P _..,.....7 • Ne.i P----1-' L \ / t °de < 440)---lif .."..f1,...„-'e ..„.7-..-1.. + t.4,6 I. n2 Vise---,-- / 0 T.,,), ?frar 't ------ik• 33? Z u*,r--.7 IE ..---"K ,... 0 17, ) < (,* ? 4 ._... / 0 t/ w I- ) \ ✓j b liA0 umue

POOM,10 9

marine units (Imperial Formation)

Figure 1.5. Geologic map of Fish Creek-Vallecito (FCV) section. - Stratigraphic symbols are explained in Figure 1.7. 10 is critical for understanding regional stratigraphy of the Salton

Trough.

The strategy employed by the present study was to concentrate field work on the FCV section, relying heavily on previous work by

Woodard (1963) and Kerr (1982), but remapping and redefining units where necessary to clarify stratigraphic relationships. Thin-section

petrography and field observations were used to differentiate Colorado

River-derived sediment from more locally derived sediment (Chapter

2). The basal nonmarine section was summarized using previous studies

supplemented by field reconnaissance (Chapter 3). Detailed measured-

section logs were prepared and paleocurrent measurements were made for

the marine units (Chapter 4) and the deltaic part of the overlying

nonmarine section (Chapter 5). Once a physical and genetic stratigra-

phic framework was established for the FCV section (Chapter 6), similar

stratigraphic units could then be recognized in other parts of the

Salton Trough on the basis of field reconnaissance and literature synthesis provided compilation (Chapter 7). The resulting stratigraphie

a basis for paleogeographic mapping and historical summary (Chapter

8).

Regional setting

Salton Trough a large non- As a physiographic feature, the Salton Trough is the Peninsular marine depression, much of it below sea level, between the Gulf of Ranges and the Mojave Desert ranges. It is separated from a minimum divide California by the Colorado delta plain, which has 1 1 elevation of 12-14 m (the topographic "spill point" of the Salton Trough).

The center of the topographic depression is now occupied by the Salton Sea. This body of saline water resulted from accidental diver-

sion of the Colorado River toward the northwest from 1905 to 1907

(discussed later) and is now maintained against evaporation by irrigation runoff (Hely, Hughes, and Irelan 1966). During much of the late 19th century it was a dry lake bed with a minimum elevation of -84 m;

the lowest part was mined for salt from 1880 to 1890 (Sykes 1937). How-

ever, fresh strandline features up to 14 m above sea level and young,

undeformed fossiliferous lake beds around the Salton Trough attest to

the existence of a large lake ("Lake Cahuilla") which occupied the

Salton Trough from A.D. 900 to 1400 and at times filled the Salton

Trough essentially to the spill point (Weide 1976, Wilke 1978, Waters

1983). Late Pleistocene shoreline traces occur as high as 150 m above sea level, presumably due to tectonic uplift (Stanley 1962, 1965,

Thomas 1963). Prehistoric fluctuations of lake level were probably caused by natural diversions of the Colorado River, alternately toward the Salton Trough and the Gulf of California.

Like the Gulf of California, the Salton Trough and the Colorado delta straddle the transtensional plate boundary between the North

American and Pacific plates. As a tectonic feature, the Salton Trough is essentially a northward extension of the Gulf of California which has been filled with terrigenous sediment from the Colorado River and more local sources. In the Salton Trough region, the plate boundary is 12 not a discrete strike-slip fault, but rather a broad diffuse zone of complex wrench tectonics (Plate 2), including (1) several major northwest-southeast, dextral strike-slip faults; (2) antithetic,

northeast-southwest, sinistral strike-slip faults; (3) transpressive folds with predominantly east-west axes; (4) crustal spreading centers and pull-apart basins which offset major strike-slip faults; (5) domains of clockwise block rotation; and (6) low-angle, normal, detach-

ment faulting (Schultejann 1984, Isaac, Rockwell, and Gastil 1986).

In contrast to the Gulf of California, where transform faults

are offset by pull-apart basins with strong bathymetric expression

(Henyey and Bischoff 1973, Bischoff and Henyey 1974, Niemetz and

Bischoff 1981), in the Salton Trough major strike-slip faults are

offset by spreading centers with little or no surface expression be-

cause of rapid sediment influx. Salton Trough spreading centers are

defined primarily by concentrated microseismicity (Keller 1979a, Elders

1979, Johnson and Hill 1982, Fuis and Kohler 1984) and are charac-

terized by high heat-flow anomalies (Elders et al. 1972, Elders 1979) and local volcanism (Elders et al. 1972, Robinson, Elders, and Muffler

1976, Korsch 1979). Studies of deep geothermal wells have documented extensive hydrothermal alteration (including greenschist-grade metamorphism as shallow as 1 km below the surface) of young (Plio-Pleistocene) terrigenous sediment, and intrusion by mafic dikes (Muffler and White

1969, Elders 1979, McDowell and Elders 1979). Refraction surveys have revealed the presence of a transitional velocity structure between the low-velocity sedimentary fill of the trough and mafic basement (Fuis 13 and Kohler 1984). In the emerging model of crustal genesis, injection of basaltic magmas in young, brine-saturated terrigenous sediment results in hydrothermal metamorphism and a gradational crustal structure at Salton Trough spreading centers.

Transpressive uplift of Salton Trough basin-fill has created exposures of Neogene and early Pleistocene strata in various locations on the western and northern flanks of the trough (Figure 1.2; Plate

3). In most of these areas the strata are deformed by folding, normal faulting, or strike-slip faulting; paleomagnetic data indicate local tectonic rotations as well (Mace 1981, Johnson et al. 1983, Bogen and

Seeber 1986). Most such uplifted and deformed areas are closely associated with major active strike-slip faults (Plate 2).

Subsurface information on Neogene stratigraphy is sparse. Wells drilled in the western part of the Salton Trough were summarized by

Dibblee (1984); Proctor (1968) briefly discussed the Texas Edom-1 well in the northern Trough. Several wells have been drilled in the Yuma

Desert in the southeastern part of the trough (Figure 1.1, Fortuna and

San Luis Basins; Figure 1.2). These wells provide the only Neogene stratigraphic data for that area (Olmsted, Loeltz, and Irelan 1973,

Eberly and Stanley 1978, Winterer 1975).

Geodetic trilateration studies have shown that transform-boundary shear strain is accommodated on short time scales (-10 a) by distributed shear across the entire trough (Savage et al. 1979, Savage slip 1983). Only about 40-45 mm/a of the total 56+3 mm/a plate-boundary has been accounted for by historic shear strain across the Trough 14

(Savage 1983). On geologic time scales, most of the shear strain has been accommodated by strike slip distributed over several faults

Crowell and Ramirez 1979, Rockwell and Sylvester 1979). Various models of "instantaneous" kinematics of southern California, including the Salton Trough region, have recently been proposed by Bird and Rosen- stock (1984), Rundle (1986), and Weldon and Humphreys (1986).

Gulf of California

The Gulf of California (Plate 1) is a long, narrow embayment of

the Pacific Ocean, flanked by rugged topography and characterized by

irregular, high-relief bathymetry. Numerous bathymetric depressions

along the Gulf axis are interpreted as spreading centers or pull-apart

basins which offset strike-slip faults of the Gulf of California trans-

form plate boundary (Larson 1972, Henyey and Bischoff 1973, Bischoff

and Henyey 1974, Niemetz and Bischoff 1981). Recent studies of the Gulf

have focused on hydrothermal systems and mineralization associated with al. such pull-apart basins, particularly the Guaymas Basin (Gieskes et

1982, Stout and Campbell 1983).

Physical oceanography of the Gulf was summarized by Roden in surface waters (1964). The highest salinities and temperatures occur the summer (Roden at the northern part of the Gulf, especially during in excess of runoff and 1964, Byrne and Emery 1960), due to evaporation Mediterranean-type, slightly hyper- precipitation into the Gulf. This the Colorado River in its saline regime probably did not prevail when in the next sec- natural state discharged into the Gulf, as discussed

t 15

Tide range in the Gulf is greater in the north than in the south, due to amplification by the shape of the Gulf. Tides at the head

of the Gulf are semidiurnal, with a mean range of 7 a, a mean spring

range of -10 m, and a mean neap range of 2 m (Roden 1964). Circulation

in the Gulf north of Isla Tiburcin is dominated by tidal currents, which

may reach 1 to 3 m/sec, accompanied by strong rips and eddies (Roden

1964). Tides at the head of the Gulf have a significant counterclock-

wise rotational component superimposed on the oscillatory component,

with the effect that suspended sediment from the Colorado River is

preferentially concentrated on the western side of the upper Gulf

(Thompson 1968). The Gulf was probably considerably narrower in the

geologic past when transtensional opening had not proceeded as far

(Chapter 8); this may have had a significant effect on tidal amplifica-

tion and maximum tide range at the head of the Gulf.

The shape and bathymetric configuration of the Gulf, combined a with demonstrated right-lateral offset on the San Andreas fault, led was formed by number of early investigators to postulate that the Gulf North America northwesterly translation of Baja California away from

(Shepard 1950, Carey 1958, Hamilton 1961, Rusnak and Fisher 1964). This the hypothesis was confirmed by geophysical data that demonstrated that anomalies, East Pacific Rise, flanked by symmetric linear magnetic and Smith 1968, extends inside the mouth of the Gulf (Larson, Menard, anomalies Moore and Buffington 1968, Larson 1972). These magnetic last -4 Ma (Larson constrain the spreading history of the Gulf for the obscure and 1972, Moore 1973); the earlier history is more 16 controversial (Ness et al. 1981, Mammerickx and Klitgord 1982, Moore and Currey 1982). Karig and Jensky (1973) reviewed evidence for a middle to late Miocene marine Gulf of California, and postulated that this basin ("proto—Gulf") was created by continental rifting possibly

related to a late stage of arc volcanism in western Mexico. In con-

trast, Moore and Curray (1982) and Currey and Moore (1984) argued on

the basis of stratigraphy and geophysical data at the Gulf mouth that

the Gulf was initiated at 5.5 Ma by translation of Baja California at

an essentially constant rate. However, Smith et al. (1985) recently

confirmed earlier disputed reports (G6mez 1981, Stump 1981) that marine

waters were well inside the Gulf by 12 Ma (middle Miocene).

Paleontologists have remarked on the similarity of the Neogene

invertebrate macrofauna of the Salton Trough to tropical Caribbean

faunas as opposed to temperate Pacific faunas of the California coast

(Arnold 1904, Vaughan 1904, Stump and Stump 1972). This similarity has or been attributed to recruitment from the Caribbean during the Miocene now early Pliocene prior to emergence of the Panamanian land bridge,

thought to have occurred during the Pliocene (Keigwin 1978). Preserved and are Neogene macrofaunas of the circum—Gulf are mostly calcitic and echinoids (Durham dominated by Ostrea, Anomie, pectinids, corals,

1950, Smith et al. 1985). sedimentation Einsele and Niemetz (1982) estimated average Gulf of California. rates over the life of the central and southern which 0.025 mm/a was They estimated a total rate of 0.036 mm/a, of an average attributed to terrigenous influx; they also estimated 17 denudation rate of ^10.02 mm/a in drainage basins discharging into the central and southern Gulf. In contrast, Johnson et al. (1983) estimated Plio-Pleistocene sedimentation rates at the FCV section in the Salton Trough ranging from 0.5 to 5 mm/a, or approximately two orders of magnitude higher. Sedimentation rates in the Gulf of California are locally high as well; Einsele and Niemetz estimated at least 1 mm/a for the Guaymas Basin.

Colorado River and delta

The Colorado River can be characterized as a scaled-down ver-

sion of the Nile; the Colorado is about one-fourth the size of the Nile in discharge, length, and size of the irrigable delta plain (Kniffen

1932). Like the Nile, the Colorado crosses an extensive desert and

supports agriculture on its delta plain in an otherwise hostile cli-

mate. Because of human modification during the 20th century, the Color-

ado River that now enters the Salton Trough and discharges into the

Gulf of California is much different from the river in its natural

state. For this reason, early studies by Kniffen (1931, 1932) and Sykes

(1937) are important for inferring the natural regimen of the river.

The most drastic alterations of the river resulted from (1) the 1901

diversion of part of the river near Yuma for irrigation; (2) accidental diversion during a flood in 1905 that rerouted the river northward, creating the Salton Sea in the process; and (3) completion of Hoover

Dam in 1935, which created Lake Mead; this reservoir traps most of the sediment from the upper Colorado River and regulates downstream discharge. 18 The Colorado River drains an area of -625,000 ke, including parts of six states. Records of the Yuma gauging station during a 24-year period before construction of Hoover Dam (Kniffen 1932) provide the best estimate of discharge into the Salton Trough before human

intervention. During that time, the mean annual runoff was 2x1010 nP

(17,000,000 acre-ft). The typical annual hydrograph was strongly seasonal, with a fairly constant discharge of -370 nP/sec (13,000 cfs) from

October through March, increasing in the late spring and early summer due to snowmelt, reaching an average maximum of 2300 mP/sec (81,000

cfs) late in June, and declining to a minimum in September.

Occasional flash floods from various tributaries occurred at other

times of the year. The maximum and minimum recorded discharges were

5300 m3/sec (186,000 cfs) in June, 1921, and 34 m2/sec (120 cfs) in

September, 1924. Before completion of Hoover Dam, the delta plain was

flooded annually, which at one time supported a large agricultural community of Cocopas Indians (Kniffen 1931). After Hoover Dam was

completed, mean discharge of the Colorado River at Yuma declined from

-600 e/sec to -120 mP/sec (Roden 1964).

Sediment load of the Colorado has been estimated by standard hydrologic methods at the Yuma gauging station (Kniffen 1932) and the

Grand Canyon (Judson and Ritter 1964) and from filling of Lake Mead after completion of Hoover Dam (Thomas, Gould, and Langbein 1960). In

1904, a year of below-average discharge, the sediment load was esti-

mated to be >1.1 x 10 " kg, but Kniffen (1932) speculated that the average annual load was closer to 1.45 x 1011 kg. Thomas et al. (1960) 19 estimated that 1.8 x 1012 kg of sediment accumulated in Lake Mead over a 14-year period, for an average rate of 1.3 x 1022 kg/a. Judson and Ritter (1964) estimated an average annual sediment load of 1.6 x 1011 kg in the Grand Canyon, or about half that of the Mississippi River. These estimated sediment loads translate to sedimentary volume of 0.06-0.08 km3/a (assuming a grain density of 2.7 and average porosity of 30%), sufficient to fill an area 100 km x 100 km to a depth of 6-8 km in one million years.

As a result of dams, diversions, and irrigation, relatively

little fresh water or sediment now reach the Gulf of California. When

the river was in its natural state, fresh, potable water could be

obtained at the estuary mouth during flood stage, while during low flaw

brackish water could intrude as far as Mayor, some 60 km upstream

(Kniffen 1932). Since these modifications were made, the estuary has

experienced permanent saline intrusion and bank erosion. These changes

must be considered when drawing analogies between present conditions

and the geologic past. Salinity distribution, sedimentology, and ecol-

ogy of the modern Colorado estuary and northern Gulf of California may

be representative of times when the river discharged northward into the

Salton Trough, but are certainly not representative of times when the

river discharged entirely into the Gulf.

The Colorado River enters the Salton Trough through a gap between the Chocolate and Laguna Mountains; the smaller Gila River

enters between the Laguna and Gila Mountains and then joins the Color- ado (Figure 1.2). The combined river flows to the southwest down a 20 low-gradient (-0.05%) alluvial plain; repeated avulsions have created a radial-bipolar pattern of distributary meanderbelts, with a Gulf flank to the southeast and a Salton Trough flank to the northwest. This alluvial plain creates a divide between the Salton Trough and the Gulf of California, with a minimum elevation of -12 m.

The delta plain is asymmetric from southeast to northwest across the divide (Kniffen 1932), with a steeper gradient on the Salton

Trough flank (-.08%) than on the Gulf of California flank (-.035%)

(Elders 1979). Gilmore and Castle (1983) suggested that the divide is maintained in an asymmetric position by tectonic uplift, as inferred

from geodetic releveling data. Since the delta plain is a region of net subsidence, however, the reported uplift, if true, could only be a transient phenomenon. An alternative explanation for the asymmetry is that when the River discharged toward the northwest in prehistoric times the lake level was considerably elevated, so that the northwestern slope of the modern delta plain was deposited primarily as a lake bottom or prodelta rather than as an alluvial floodplain.

Prior to channelization for irrigation, distributaries on the delta plain were characterized by meandering channels, although deviations from classic meanderbelt patterns were common, including anastamosing channels and local straight reaches (Sykes 1937b). Drastic changes in the main Colorado channel occurred between surveys in 1858 and 1891 (Sykes 1937b, fig. 15) and in the estuary between 1826 and

1851 (Sykes 1937b, figs. 5 and 6). At the Yuma gauging station before completion of Hoover Dam, the river bed underwent severe scouring 21 during floods and backfilling during low flow. Kniffen (1932) illustrated nearly 15 m of river-bed aggradation between the peak discharge of

June, 1921 and the low flow of September, 1924.

Van de Kamp (1973) studied cores taken along the New River

distributary and presented a generalized profile for a channel-fill

sequence; he depicted a fairly typical upward-fining point-bar

profile.Stephens and Gorsline (1975) described the sedimentology and

stratigraphy of the New River delta formed in the Salton Sea since the

1905 diversion.

The marine flank of the Colorado delta is of particular inter-

est because it is a well documented example of a modern tide-dominated

delta in the classification of Galloway (1975). Sedimentology of this

area has been studied by Thompson (1968, 1975), who concentrated on the

tidal flats, and Meckel (1975), who emphasized broader facies relation-

ships, vertical sequences, and sand-body geometry of several deposi-

tional environments. Meckel (1975) emphasized the presence of cur- the river rent-oriented submarine to intertidal ridges offshore from estuaries mouth; such ridges are typical of tide-dominated deltas and of the modern (Off 1963, Wright, Coleman, and Thom 1975). Tidal flats and evaporites; disruption of delta are dominated by deposition of mud minerals is an important sedimentary structures by growth of evaporite

process (Thompson 1968).

Effects of human modification of the Colorado River on the reports. Kniffen marine delta are difficult to assess, even from early he reported (1932) visited the deltaic coast in 1927 and 1928; 22 difficulty in landing due to the shallow, muddy bottom and extreme tidal range, and Observed that the tidal flats were coated with thick,

liquid mud after each recession of the tide. He also reported that

Colorado River alluvium extends southward along the coast of Baja

California nearly to San Felipe. Kniffen (1931) reported large shell

middens of the Cocopas Indians along the shoreline, and stated that

species of clams in the middens were similar to those of the modern

coastline. These middens could potentially be useful in interpreting

paleoecology of the delta in a natural state, but Kniffen did not list

species or give other details.

One effect of the high tide range is tidal bores in the Color-

ado estuary. Such bores were typically -1.3 m high on spring tide, but

could attain 3 m in shallow water (Kniffen 1932) and were a significant effects hazard to navigation (Sykes 1937). However, the sedimentologic

of tidal bores in the Colorado estuary have not been studied. system in Origins and early history of the upper Colorado River controver- the Colorado Plateau and Grand Canyon have been a subject of

sy for several decades (e. g. Blackwelder 1934, Longwell 1946, 1954, River downstream Hunt 1969, 1976). However, the history of the Colorado to subsurface from the Grand Canyon is fairly well established due Project (LCRP), primari- investigations during the Lower Colorado River and ly an investigation of hydrologic resources (Metzger, Loeltz, and Irelan Irelan 1973, Metzger and Loeltz 1973, Olmsted, Loeltz, and terraces are 1973). Above Yuma, the lower Colorado river floodplain Bouse Formation (Metzger incised into marine beds of the ?Pliocene 23 1968, Smith 1970, Metzger et al. 1973, Metzger and Loeltz 1973). The few wells that have penetrated the entire Bouse section beneath the Colorado floodplain have encountered locally-derived sands and gravels below the Bouse, but no Colorado River sediment. Therefore, the age of the lower Colorado River is probably post-Bouse. Shafiqullah et al. (1980) reported K-Ar dates of 5.5 Ma from an ash bed at the base of the Bouse Formation, and a K-Ar date of 3.8 Ma from a basalt flow overlying a Colorado River terrace within 110 m of the modern flood-

plain. These dates thus bracket initiation of the lower Colorado River.

Climate

Except for Death Valley, the Salton Trough is the most arid

region in the conterminous United States. Average annual rainfall is

<10 cm over most of the Trough (Pyke 1972), while annual evaporation

from the Salton Sea has been estimated as 1.8-2.3 m (Hely et al. 1966,

Weide 1976). Recorded temperatures at Borrego Springs range between extremes of -9°C (15°F) and 49°C (121°F), with a mean of 21.0°C (69.8°

F) (Felton 1965). Daily temperature fluctuations are large, especially in winter; a daily range of 20°C (35°F) is not uncommon. Prevailing winds are from the west in San Gorgonio Pass, where windmill farms are used to generate electricity. The northerly Santa Ana winds, associated with winter storms, sometimes affect the Salton Trough (Bailey

1966). On the lower Colorado delta near the Gulf of California, prevailing winds are seasonal, changing from southeasterly in the summer to northwesterly with frequent gales in the winter (Kniffen 1932). 24

However, the northern Gulf lacks constant surf, and most beach erosion occurs during storms (Kniffen 1932).

Pyke (1972) documented two seasonal precipitation maxima in the

Salton Trough. A December maximum, primary in the western Trough but secondary in the eastern Trough, is associated with extratropical

Pacific cyclones. An August maximum, primary in the eastern trough, is

associated with convective thunderstorms. In Baja California, early

fall precipitation results from tropical cyclones, some reaching hur-

ricane intensity. According to Pyke, surface temperatures of the Paci-

fic Ocean are the dominant influence on precipitation patterns in the

southwestern United States.

Little is known of Neogene and Pleistocene climates in the

Salton Trough. Sparse evidence for more temperate pre-Holocene climates

will be reviewed in Chapter 8.

Basement geology and Flanks of the Salton Trough consist primarily of Paleozoic

Mesozoic igneous and metamorphic rocks with minor Tertiary sedimentary on a and volcanic rocks. The following discussion is based largely

recent review by Sylvester and Bonkowski (1979). Ranges consti- To the west of the Salton Trough, the Peninsular intruded by the Penin- tute a west-tilted block of metamorphic rocks the Santa Rosa Cataclas- sular Ranges Batholith; these rocks are cut by attain greenschist to amphibo- ite Zone (SRCZ). The metamorphic rocks and lesser vol- lite grade; the protolith was interbedded sediments in age. Miller and canics, generally regarded as Paleozoic to Jurassic 25

Dockum (1983) obtained Ordovician conodonts from marble in the Coyote

Mountains. The Peninsular Ranges batholith consists of calc-alkaline granitoid plutons of early to middle Cretaceous age. Tonalite is the average composition and dominant lithology, although a range of compositions from granite to diorite occurs in the batholith. The SRCZ consists of rocks ranging from augen gneiss to ultramylonite; it cuts

the plutonic rocks and is overlain locally by Eocene conglomerates, which constrains the age as late Cretaceous to Paleocene. The SCRZ

provides important piercing points for determining the offset of the

San Jacinto fault (Sharp 1967). Simpson (1984) inferred west-directed

thrusting for the SRCZ, whereas Engel and Schultejann (1984) inferred

north-northwest tectonic transport. Engel and Schultejann (1984) des-

cribed a complex post-batholithic history for the eastern edge of the

Peninsular Ranges, including late Cretaceous folding and imbricate

thrusting, and mid-Tertiary detachment faulting.

Three ranges with basement exposures, the Vallecito, Fish (Figure Creek, and Coyote Mountains, occur adjacent to the FCV section plutonic 1.3, Plates 3, 4). The Vallecito Mountains consist entirely of local diorite rocks; Kerr (1982) described them as quartz diorite with aphte, and quartz monzonite, intruded by dikes of "spotted" tonalite, except that bio- and pegmatite. The Fish Creek Mountains are similar, consist tite schist occurs locally. In contrast, the Coyote Mountains Marble is the dominant largely of greenschist-grade metamorphic rocks. including Carrizo Moun- lithology in the easternmost Coyote Mountains, in the tain (Miller and Dockum, 1983), while biotite schist is dominant 26 central and western Coyotes. Other lithologies occur locally in the Coyotes, including amphibolite, gneiss, quartzite, and intrusions of quartz diorite, granodiorite, and pegmatite (Christensen 1957, Ruisaard 1979). The SRCZ does not crop out adjacent to the FCV section. Basement rocks along the eastern margin of the Salton Trough

are very heterogeneous and their history is largely unrelated to that

of the Pensinular Ranges. Local occurrences of unique rock types have

been important in establishing offset on the San Andreas fault system;

for a review of these rocks the reader is referred to Crowell (1975) and Sylvester and Bonkowski (1979).

Previous work on Neogene stratigraphy of the western Salton Trough

Neogene strata crop out extensively along the western margin of

the Salton Trough (Figure 1.2; Plate 3), where they have been studied

for more than a century. The Neogene stratigraphy of this region can be

concisely summarized as a marine section bracketed by nonmarine units

(Figures 1.5, 1.6, 1.7). Stratigraphic nomenclature was more or less

standardized by Dibblee (1954), but usage varies even among later

publications. A significant complication is that stratigraphic se-

quences can vary considerably over short distances due to facies

changes and pinch-outs. Therefore, it is essential to separate problems

of synonymy within a given section from problems of regional correla-

tion; this separation has generally not been made in the past. Synonymy

tables for the western Salton Trough are shown in Figures 1.8-1.11; similar tables for other areas are included in Chapter 7.

•

14 0 • - cd 0 .1-1 0 4) 0 4-) 0 .ur1:3 1711 C+ri g >1 g Cd 8 2 g 44 S 2 2 O 0 fl, c0 ,- 0 C..) $.4 0 CO +) 0 = .0 e-) 0 P.) - Z 0 T) 0 04 CD Z 4.) '''l , 7 0 ,0 0 go +) .1-i in r. a' g. 8 -;:', P'41.3 ' ai. 0 go X 0 fal in ... in 0 Ci/ V 0 4) t) • - >.) ci-4 0 0 14 -,-) a.) - i - o xo 44 .0 'DO A. 114 1.4 V) r.) E .0 -R 1 •P 0 0 cil *) C) C.) cd()4clula).-841 0 04 ..4 ..0 i., 0 = El rml 1-1 ei +) 41 gg 0 -0 H 03 4) >'s - .. ai • r-4 0 o V g 0 co .1-1 CN1 O 4-) • 1-4 Z 0 CD E 13:1 1:::1 ! N VI 0 1:4 C.P V •r4 a. g " 1.4 • • n•.1 •• a) ii"o.. -0 i.. to cp., cn ix .,g ci i. cl 1 fo o 1=. 4 9-1 o 01. c ,--4 c .0 o - u P - P.) 0P.) E 0 0 ,--1 g g4g g 0 N 0 0 ?),-4 O 0 0 go P.) vo 0 02 0 •ri .0 a 0) 0 op ,-i *) C.) C.) C) c.) C.) --0 CD CO z - In 0 C'.; 0 .1 cf) c. N - 0 ,-i C.) 0 CD X - 0 • > TI .. ^0 0 ;r.4 4.) •--I L..) W a) • •• g c 4,01a.)04) "C' - -0 ca c .1-4 0 P. 0 f() -•• 4) 44 +) 0 P-. 4) .0 I 1. O ;:l S 14 I I V 6 4 E I ..4 0 C.) -0.m-.4 CO -co c/‘ o a) c +) -4 u2 c ,-.4 -0 14q 0 0 0 • n-i Ca ....ri ixi 14 ••. .P4 •ri ••i 1.1 ••• 0 .0 0 ..0 -0 g 0 '0 0) C.) 04 C.) al g el -,-. ..a) 4 I.' 1) 5 c.) 000ç X .0 O ii - C.) C.> 0 u) 4) -,-1 .1-.1 C.) 0 ,.-i cc-( -0 "0 4t ..-1 ,--4 0 g E g 42 0 0 1-• O 4-1 i-f 0 0 0 0 •,-4 0 -P: 0 0 0 @ i-i. 0 0 0 t- P.) - 0 0 g B 0 • 0 C,) C.) ›) 0 0 (1) • r4 ,•••n op ..-4 -0 -4-) C.) 0 i 0 ...d' -P C.) C.) a) - 4-) 0 c. 0 tr3 p $,--1 0 0 0 2 0 ,... p .•-1 o &., si 4, o ci -1-) tic GI) >, •n-i 0 0 to •,-1 0 = ct., .4.4 a .x -1- • 0 4., o 0 C.) Q) Q) r0 fr., 0 - ›., 0 0 (11 U g C.) 1-4 Z 1.4 (1) a) .44 ... •P 0.. C.) •O 14 • Q) •• .. 4.) 0 Ca 10 Cr.1 0 . - >.. lx 4 'DO qt) g t4 Z rp •n•• •r-I p ... .ii. ji .1-1 -6. rz4 CC .r4 g 4-) 0 0 oz 4-) ii-) ii''' 1 g ,--, >, c to 110 - -1-1 Ô P' P5-pe °0 0 0'A 6 c?, z 4)0MZCIC/ L7.4 E-) 28 29

BORREGO BADLANDS & SAN FELIPE HILLS >-fcc colluvlum Qc 1Qall alluvium Brawley2 CC Qb Formation Qt terrace gravels z Ocotillo' o Qo Conglomerate Hues& Canebrake' Pah member } Palm Pac Conglomerate Pb Borrego' Formation UPPER Spring NON MARINE Tap'ado' Pt member UNITS Formation 01la Diablo DELTAIC SEQUENCE Po member Pd member Jackson Fork Camels Head member member Deguynos and Pj Pch Pdcu Camels Head members (undif I erentiated) Deguynos Pdg member Lavender Canyon Plc member Mud Hills Pm member Wind Caves MARINE Pw member Imperial UNITS MPiu Formation upper Stone Wash (undifferentiated) ub boulder bed MPs member upper MPlu turbidite lower Lyclum Kamla Creek Latrania M P I I turbidite MPI member MPk member MPla member MPIt transitional

Mf Fish Creek Gypsum

oLuzwi upper Meu f anglomerate- sandstone tholeiitic I LOWER o Elephant basalts Alverson lower Mat NON MARINE Trees boulder bed Me 1 Split Formation UNITS lb member alkalIc Ma Mountain Maa basalts lower Formation Mel fanglomerate Red Rock Mr member crystalline basement, undifferentiated (Cretaceous plutonics & high-grade metamorphics with Ordovician-Jurassic(?) protoliths)

'Ocotillo Cg. Is lithologically equivalent to Canebrake Cg. Hueso mbr. 2LacustrIne units.

in this report. Figure 1.7. Stratigraphic units and symbols used

30 -0 CO LO ID 0 -0 '',.ANNA.A.MAA, .4* Tt Ch CV v- CI3 ).. CO -0 a • tili co .6 (idN) %qui 2 'r...... ren .2 ET; ..2.i. mee.so alun)! C 0 41 .) ° E .-- c ô.' • a: - c — — s 'Fu f, 1= e. NNOAAANNAN• 0 co E o. E p.., o. >•.-- t 5. ,-- 0 c IX ., = . _, (...) -,,, ,.: ,_46 0 a t Q) ii), — g:6 Alvo,w, -0 6 o .0 (seiciluVue) J;Juclitu zra5.) Ifi il, E (-) X o E 2 E .,." 8 lud &los wied uclisuuod ispeduil u- limules

ccr 8 (D 8 — 8 c .6. 0. 8 .0 c E c0 aci E -0 E 8, -D a) E 52 0' D -ff ? .o 00 .- E uouewiod lepeclug C 0 sape' 9 I V u) a <7; co 2 Seipel Sepe; g e) S , E .. E c .r. 0 — ... •.a spues ,_ 0 _. = t. o2 Tu ',7 '" sAsio u!e4unoiry e4oAop 0 0 e!usilei > -0 Z ai ro 8 Q) 0 u. uoweiwoj lepaciwi Zr < Tr m) c „,0 /..ANAT; E m n 73 g o '=. 0 m -1, .3 0 : 0) 7. ...-: 8 7 a) a, 2 .'@. f-E721-2-Fg o. E ; E 2 — E 6 a- — Li. .:' 2 ....d ozuv .-ii . . . C 70 g 0 0 'ai a) 0 -..": = ...-Ta •-• -2 m N 'al œ .-eo —c0 0 0) o. E g,' E c?)- E 6 u-8 2 u_8 c .T. 6 c c .c g 0) .2 à , % " g 2 .- Fa — E .c 7,3 Jeqwew .pueqo.uns ,_ E ,- 70 4 '6 7s Tu 6. E 713" a' 4"E ..zE 2°' 0.. co 8 _, E .2 ,- n CO ..- < o CO o -• la_ uonBun0J lepadwi u_ 2 2 c g -,2 .g 4-) C-0 o ._ E E . Ln .r. • co o. E a> m .o >'a-) E uouecuJoj CJ) '.- .rt o uleitmovg uogew.loj igpartiM U. Pros •0 C . t *0 6, ; L_F___ c t .6 g , o 'o03 .0 ° r.,-.. g 0 a ; Go c75 .0 `c,' -0 c r... E c --;-- u E ra E g. E .1 a 0 ri, S'" 2 g, it, — i: W cr, 0 CO 0 ro ?..:' w o. E ' cp 9 'wew 'Ir' E uolteumj 0 e 0 E CD E .oPBJPu‘e > ' ure; unerj r Zr 0 c.) IL uo!luunod iBpacitm la. 4iids C C — o C 0a) Cl; 0, .2 as -- a> Tt E c 7,,, T.: ''" 8 -77; = ; 70 Zi 0.o a' E E _ c 6 0 0 E 6 < < 2 oc! u. — U. c as e c •c `c ô T0 .2 5), ô .12. = Ta, = ,- al.* ,_ ›, as 6. , ao 0, E '4" ii-) 7° a' cE moE > cl, n- - 0 .....- ,.. o .- — 0 0 m ..., o i E 8 < LL 1- — LI_ u. c 1 ci, .{:,, 2 i 0,C 2 i .0 .0 .- es, E .c (1 1 :=.- 0 E E -6 co .0 o 0 Tu 6. E I o •'- a. (,) 8 I uopeunoj lepacitm U. I ..„5 e_ I ci = I Ia ...... c73cl). '';„ 'E-ci .- c co m = .e' ° .mcr, E I I .0 0 C,),.. .C.. co c,.'. ,.„ o. C CSI m 0 c u- c- al C73 ..S i7i 0 2, 5 7,6' E o I >- cc x '- 0 m _j i.."-An., 0 I a I uo!s!A a uoisliqa I ô., Jaddn lemoi 0- I i ce I •spaq.Naela oz! I JO 'WJ PZPJBO . I To 0 . l' E 0 (if C c , > C O a) 2 2 0 7-7 23 Tc 0 010 .0 0 0 .0 0 C .- 7 g 8 w m - e' z ts

n. -0 0 .0 co U) a v nt nt t\I •-

>. D (no) Jeddn (Ism Joonoi 0 Si = E - i e.-. 2 3 -0 o ,.: 0 F, 9-- 8 . . o. ± ''' '' it 0 , -.; 2 1Z - -0 ô (-) ,.; n- o E .0 o .0 a D ." x E E o E o E Y .0. 8 T. - 0 :6, 1: ccr. '5'- I - ' .5 .g 0 0 ..S e 0 t t -0 ••• 0 E V 0 w 0 e.1 uoliewaoA BupdS u_ 0 wled uoueLuJoA lepocluit 'Ind u!ewnoiry picIs c 61 f I 0, 2 Jecidn E c •-• I eIPPIw = 17,7 . To 7- c- 52 0 a. a :2• S E W I o .- 0 (-) uonewiod lepoduil 1 M 2

1! Z I m = = 2 ..2 1 u .. co < 2 1 8 8 8 1 cc cc cc ... 0 . c, 0 2 2 2 -8 1.5 ---°0 -E (-) -3 20 ow 0 .„, al c, ... ,_ u- ô ô ô co ô .3 -..F. .._ 0 0 O. - = = = = ... G3 U . t g = in : Fj f, 2 eoppicuni vd,! scollop:pm 2 ,,2 a 4,,, a ... = .6- ' ° & % g -f.'- .1=' 13 7! P- 0 uolieu.uod leiieCluii U0UULI0j EVA, cl w. ô ô 8 ,: .0 .0 •: gi' ..: t o ci., n: o. .lo 3 o. -0 3 .0 - : "a .o a E • E 0 E a E I E 0 E m 1, E ôtB . E - 0 _ E co gt 0 .-. OD O -Wd BWAS WIEd U04eunoj lepadwi •tud u!elunoVi 4! 10S •LUd eZUV c d. 0, 0 c, E c 47. Jeqwow puagonne Jaqwow e!ue.nei u. -0 ,_,,_ = E E t•••• 73 7: w 0,1,.t 7 CD 0. a E co co 8 uo!letuioA reliackui •wA wewnotry ivls 0 • ..npuqns uqums • -o • c,,,,,...... „ Joddn Ô te 71) -2 .. Jeden Jemoi 7 .:• o . ô ,: Ô ., .0 (0 . 4,.."94, • JemOt g 1 t >, E !..._ g ,_, g 2. E o E 8 e2 x uqui mum L) •iqui sowinBoa uo!cow.,od 6wids wied uoilewiod lepodwi •toj welunorry pids .u.u.ezuv c 2.. Td ..: .: f.. 2 IF, "0 o 0 .0 c- .o. -2 .o a E 7';' E .o!,, E E ,- E 1 œ TO a Tc- .°1E o E o O. co 8 w uonew.loA iepedwi uo!lew 0A wewnoini olds co 0. .0 ....,4> = 0 c .10 g I '?'2' i (:)T :20 ..`:E c.0 f 0 2 o > .3 ..., to g SOUSEI0 OS.1800 E g.-- al a CO . E ; i.:6g. snc,apusso)un

0 U0!I?W.103 wewnovi mds • n c 6' eu v ioi S 2 .s.co .9.zii . 2 ô ' TO 0 a) ‘,'2 0 8. E E o .0 cp g g OVV.Ws 13' - ô uo!iewmA wewnoiry uicls V ° u_ c • ,i 713 .2 0 .8 00) E - E c u-, 7-3 c.- nJ 0, a E -sm...., Er. .- E ô -wA wetunoN tuds u_ ' a! . • c g . e . a; .z 15 E c 7,, a .r, 0 E :','. E _o S '''E o E 0a 0 8E co 8 M u_ .. g u_ uollewiod lepachui 32

Dibblee, Woodard, Kerr, Pappajohn, Richardson this study 1954 1963, 1974 1982, 1984 1980 1984

Canebrake' Canebrake Canebrake Cg. Cg. Conglomerate (Pao)

E Rock u.. • Rock Unit co 0 Diablo Palm c Unit D •- mb r. C C. 011a* 7 Spring B. cr) t mbr. (upper Rock Eu_ Fm. c submbr.) Unit To TEI 13 a (Po) EL

Camels Rock Jackson Head* Forembr. c Unit A 0 mbr. (PD "iv: c Imperial Imperial OE. ,'73 IF; upper 2 Deguynos 4c — E submbr. Formation mbr. (Pdg) Formation 70 *co z o c W c o a ), lithofacies lo E c lower lli Mud Hills* _ oi E 4b a.) submbr. F `6 mbr. (Pm) o LL

upper Pliocene 70 Stone Wasti 3a lithofacieS V. fanglom- continental A & E 2.(1) mbr. (MPs) c erate rocks Split 73 c .c.., .2 marine Pliocene lithofacieS Lycium 43 3b arenite marine mbr. a. 2 E B 0. Mountain member sandstone (with A) (MPI) D 10 0 *a, lower Pliocene c 0 iai Formation fanglom- ? continental - o. co 2 13 o. erate rocks alluvial 'C' 1- 2 p c E •-• — c fan o „ t . 2 o t M - 0 15 co - upper h- alluvial facies .- -0 E a) N *CI'S fan 1= Cl. c E member (43, a o o GC ,- cO LI u_o

'first use of name

Figure 1.10. Stratigraphic synonymy for Split Mountain west of Gorge and southeastern flank of Vallecito Mountains. 33

U) CD CO V C.) VOMAAMM nt d' N. N.. >, -0 y 0o o2 co Co' c ----0 2 o o ... ta x a, c a. :H ,,,,, >,•-• >.. c w E :1.-.. ti • 3 • ca o oc 3 E -E 0 0 0) a a) :*--- el E 0 E -g .171 .- ca nmAneVW,w,, (") œl E ô E Cod) 'iqw 02 8 8 'ô *Buo LL u_ (op) AftMmemee .60 (°od) .60 oml000 ame.sqaueo C od o , sec e o ..r..2 0 .,,, .0 c 'E '4. E .c o 3- - E c o 2 Tél `c5. E =0. E .— CL co ,,-, 0 ° LZ 'w j lepedwi

n:' ).. OC a) a> • ci) o c i•-. - E 0 N. u_ 2 "c-.; co 0) ô E CO ,-- te. co '60 al 8 oim000 u. i c .i• O 0 ._. C N- O 0) 2) cal E 0 '- 8 0 co co u_ .wd Bupds m'ad n.: ID Lo vefltd >10 Y .0 03 0 cr) o ,- El. E = a> • c 1-_,. rili (S15 E '' C -,-3 o c o) * o o c c • - -..-: cl- 0)0 E o,- o ce da ›... 7.; cd - E a> 0) o o ,_ 1-, Ci LL c. E - ,-- E ô .0• E 8 E a.- cr) .o ga: cd 8 03 " - u. ,... •60 o E ô) to u- u. pillioao eleicieueo c c * c a, o 7.,. o 4-• 0 .2 CD •-• CY) .x EC ..3 %7 fo da to f2 c° W a> ch .- E al o. E o. E o a- CL n- cd *-- (1) E ô I— cr3 0 0 Li_ LL — lt. 34

Early geologic investigations of the Salton Trough focused on the drainage basin and flanking mountains of Carrizo Wash (Plate 3), which constituted part of the Butterfield Stage route that provided the principal crossing of the Salton Trough between Yuma and Los Angeles.

Fossiliferous Neogene beds were discovered in the Carrizo Wash area

during preliminary railroad surveys sponsored by the federal government

(Blake 1853) and subsequently attracted the attention of a number of

paleontologists (e. g. Arnold 1904, 1906, Vaughan 1904, 1906, 1917,

Mendenhall 1919, Kew 1914, Hanna 1926, Woodring 1932) who recognized

strong faunal affinities with the modern Gulf of California and Carib-

bean Sea. Kew (1914) proposed the name "Carrizo Formation" or "Carrizo

Creek beds" for the fossiliferous strata, but due to prior usage of the

same name in Texas it was subsequently abandoned Hanna (1926) recog-

nized four units within the marine sequence, including the "Imperial

Formation", a "coral reef" exposed at Alverson Canyon in the Coyote

Mountains. Woodring (1932) disputed the reef interpretation and Ob-

served that the fossiliferous bed is a local lens within terrigenous

deposits. He proposed that "Imperial Formation" be extended to include

the entire marine sequence within the Salton Trough region and proposed

"Palm Spring Formation" for the overlying nonmarine sequence; this

nomenclature has been followed more or less by subsequent investigators

(Figure 1.8).

A shift of attention to the adjacent Split Mountain-Fish Creek

Wash section (Figures 1.3, 1.5, 1.9, 1.10) was prompted by acquisition

and closure by the U. S. Navy of part of the Carrizo Wash basin for use 35 as a bombing range (Carrizo Impact Area) clueing the Second World War and Korean War, by improved access to Split Mountain Gorge, and by interest in the exceptionally high-grade gypsum deposits near Split

Mountain (ver Planck 1952). Tarbet and Holman (1944) proposed "Split

Mountain Formation" for coarse terrigenous strata exposed in Split

Mountain and "Alverson Formation" for volcanic rocks below the Imperial

Formation in the Coyote Mountains. However, their definition of Split

Mountain Formation also included some marine strata, which was somewhat at odds with Woodring's (1932) proposed usage of "Imperial" for all

Neogene marine strata; this definition of "Split Mountain" has subsequently led to much confusion.

Dibblee (1954) provided the first detailed geologic nap of the

Salton Trough, which tended to stabilize nomenclature. He extended mapping northward into the San Felipe Hills and Borrego Badlands (Plate

3, Figures 1.1, 1.11), where he recognized additional units, the lacustrine Borrego and Brawley Formations and the Ocotillo Conglomerate

(Dibblee 1954, 1984).

More detailed stratigraphie subdivisions at member rank, accompanied by large-scale maps, were proposed by Christensen (1957) for the

Coyote Mountains (Figures 1.1, 1.8) and by Woodard (1963) for the Split

Mountain-Fish Creek-Vallecito Creek area (Figures 1.9, 1.10). Unfortun- ately, the genetic significance of these subdivisions was apparently not fully appreciated by most subsequent investigators, who generally rejected the more detailed subdivisions (e. g. Smith 1962, Stump 1972,

Bell-Countryman 1984). Even Woodard dropped some of his own earlier 36 (1963) subdivisions when his work was finally published (1974). A series of sedimentological studies by students at San Diego State

University (Pappajohn 1980, Ruisaard 1979, Kerr 1982, 1984, Kerr,

Pappajohn, and Peterson 1979, Bell 1980, Bell-Countryman 1984, Richard- son 1984) used more of a genetic-stratigraphic approach, using informal lithofacies designations in lieu of formal nomenclature (Figures 1.8- 1.10).

Until recently, age control was provided primarily by biostratigraphy. Marine invertebrates of the Imperial Formation have been interpreted as Miocene (Blake 1853, Arnold 1904, Mendenhall 1910, Kew

1914, Woodring 1932), Pliocene (Vaughan 1917, Hanna 1926, Durham 1960), or late Miocene to Pliocene (Stump 1972, Ingle 1974). Stump (1972) provided the most detailed biostratigraphic description., based on foraminifers in the Split Mountain-Fish Creek section. He recognized

Pliocene planktonic zones N19, N20, and N21 (Blow 1969) in the Imperial

Formation. He included the marine beds of Tarbet and Holman's (1944)

Split Mountain Formation in his Imperial Formation (Figure 1.9), but regarded their age as indeterminate due to the absence of planktonic foraminifers. Downs and White (1968) interpreted the age of the Palm

Spring Formation in the FCV section on the basis of vertebrate faunas

(Figure 1.9). They interpreted the "Layer Cake" and "Arroyo Seco" vertebrate fauna as Blancan (late Pliocene by modern usage) and the

"Vallecito Creek" fauna as Irvingtonian (early to middle Pleistocene).

More precise ages for the FCV section were recently provided by

Johnson et al. (1983) based on a sequence of magnetic polarity 37 reversals calibrated by a fission-track age. Eberly and Stanley (1978),

Ruisaard (1979), and Gjerde (1982) obtained a number of K-Ar dates on basalts of the Alverson Formation (Chapter 3); these are mostly early to middle Miocene.

Study area

In the present study different areas were examined at different

levels of detail. The most intensive study, including detailed mapping,

logging of section, sedimentological study, sandstone sampling, and

paleocurrent measurements, were made in the lower Fish Creek basin

(Figures 1.4, 1.12). Supplemental work, including reconnaissance map-

ping, logging of section, and paleocurrent measurement, was done in the

upper Fish Creek basin, Vallecito basin, and Carrizo Impact Area (CIA)

(Figure 1.3). Additional mapping was done in the vicinity of Canyon Sin

Nombre and Sweeney Pass. All of these areas were studied by photogeo-

logic napping.

Stratigraphy and paleontology of the Coyote Mountains (Figure

1.1) were studied concurrently by Susan M. Kidwell. The author visited

Kidwell's localities and incorporated some of her observations in this

study. Additional information on geology of the Coyote Mountains was

provided by Christensen (1957), Ruisaard (1979), Mace (1981), Gjerde

(1982), and Bell (1982).

Field reconnaissance was also made in the following areas

(Figure 1.1): (1) the Yuha Basin, principally the area of Yuha Buttes; of the Moon (2) the Borrego Badlands, mainly Fonts Point and Hills

Wash, Borrego Mountain, and the San Felipe Hills, mainly Shell Reef, 38 39 Tule Wash and Arroyo Salads (Plate 3); (3) the Indio Hills, including the outcrops in the western hills and at Willis Palms; (4) the Super

Creek area near Whitewater Canyon, and Garnet Hill; and (5) the Bouse

Embayment (lower Colorado River valley), including Milpitas Wash and the Riverside Mountains.

Structure

The FCV section occupies a southwest-tilted (-20-25 regional

dip) fault block between the Elsinore and Coyote Creek faults (Figure

1.1). A gravity study of FCV basin suggests that basement on the north-

ern and northeastern margins forms a ramp, whereas on the southwestern

margin sediment and basement are juxtaposed along a vertical 4-km

subsurface scarp of the Elsinore fault (Gibson et al. 1984). The non-

conformable base of section is exposed in Split Mountain (although not

within Split Mountain Gorge), and the highest part of the tilted sec-

tion is exposed along the Vallecito Creek Wash. The intervening section

is more or less continuously exposed, except where concealed by modern

washes and late Quaternary alluvial-fan terraces (e.g. North, West,

Middle, and South Mesas). Structural complications include east-west to

southeast-northwest trending anticlines and synclines, previously

described and mapped by Woodard (1963), and a large number of small

faults (Plate 4). Woodard (1963) and Kerr (1982) mapped some of the

larger faults; these generally have northeast-southwest and northwest-

southeast orientations, are high-angle, mostly normal faults, and may

have significant strike-slip components (Woodard 1963). Additional

fault complexes were recognized in this study, based on photogeologic 40 and field mapping in the Elephant Knees area and upper Lycium Wash and

Stone Wash basins (Figure 1.12, Plate 4). Additional faults, too small

to nap, were encountered while measuring sections along Fish Creek and

North Fork Washes. In and near Split Mountain Gorge, small-scale, low-angle bedding-plane thrust faults were Observed, in some cases with

thrust-ramp structures; these cause repetition of section and chaotic

interstratal deformation which must be recognized when logging section.

These thrust faults are also too small to map.

Exposures

For sedimentologic study and paleocurrent measurements, fresh

exposures are required. The finer-grained and poorly indurated parts of

the FCV exposure (claystone, siltstone, fine sandstone) form heavily

weathered badlands which are generally not adequately fresh for sedi-

mentologic study, but sufficiently fresh exposures are found along cut

banks of the larger washes, particularly Fish Creek, North Fork, Arroyo

Seco del Diablo, Arroyo Tapiado (Figures 1.3, 1.4), and some of their

tributaries. Conglomeratic units exposed in Split Mountain and along rugged topo- the southern flank of the Vallecito Mountains form more of graphy, typically with box canyons and dry waterfalls; exposures

these coarse-grained units are better but less accessible.

Logistics Desert State Most of the study area occupies the Anza-Borrego (obtainable from park head- Park (Plate 3), and a collecting permit rock samples or quarters in Borrego Springs) is required for removing 41 fossils. Some of the larger washes have been officially designated as roads; vehicular access to other areas is prohibited. Camping is allowed anywhere in the FCV part of the park, and no permit is required, but no water or facilities are available. Entry into the CIA, a former bombing range within the state park jurisdiction (Figure 1.3, Plate 3), is prohibited at all times due to the presence of live ordnance. Access to the CIA for scientific

study is permitted only with the escort of a park ranger; prior ar-

rangement must be made through park headquarters. Such permission is

generally difficult to obtain and then only for very brief excursions.

Part of the study area occupies property of United States

Gypsum, which operates a large quarry at the end of Split Mountain Road

adjacent to the state park. Permission for access to U. S. Gypsum

property can be arranged in person with the company geologist at the

quarry (there is no telephone); permission may be denied at times due to blasting.

Entry to the FCV area is possible from two directions (Figure

1.3, Plate 3). The Fish Creek drainage basin can be reached via Split

Mountain Road, which intersects State Highway 78 at Ocotillo Wells.

Access to the Vallecito drainage basin is possible from Highway S-2

(Imperial Highway) from a number of entry points; the author has found the Palm Spring turnoff to be most convenient. These two drainage basins are connected by the steep Diablo Dropoff. The dropoff is usually one-way (downhill into the Fish Creek drainage) but can be climbed by some vehicles into the Vallecito basin when the dropoff is in good 42 condition, thereby avoiding a drive of about 80 km (50 miles). The

Pinyon Dropoff, which provides one-way access to the Fish Creek basin from the northeast, is very hazardous and is not recommended. Canyon

Sin Nombre, at the west end of the Coyote Mountains, can be reached

from a turnoff on the Imperial Highway just west of the Badlands Over-

look (Figure 1.3, Plate 3).

The condition of washes for driving is highly variable, and is

determined largely by flash flooding during the late summer monsoon

season. In some years a wash might be very sandy and require four-wheel

drive (rear-engine or front-wheel-drive drive vehicles are frequently

satisfactory); in other years the same wash might be very rocky, requi-

ring high ground clearance. Some washes marked as "jeep trails" on USGS

topographic maps are now inaccessible or have been closed by the

park. Driveable washes at the time of the study are indicated in

Figures 1.3 and 1.4 and Plates 3 and 4. Current information about person (there is no driving conditions in the washes can be obtained in ranger station telephone) from the park ranger at the Elephant Trees due to extreme (Figure 1.3). Work during the summer is not recommended months from January temperatures and the danger of flash flooding; the

to April are recommended for field work.

Methods

Mapping Woodard (1963), and Geologic maps at 1:24,000 by Smith (1962), field relation- Kerr (1982) provided a good overview of stratigraphic did not adequately portray some ships in the FCV-CIA area but 43 important details. Therefore, remapping was undertaken of Split Moun- tain and the SE flank of the Vallecito Mountains in addition to reconnaissance mapping of the Fish Creek and Vallecito drainages and the

CIA. Unpublished mapping by park ranger Paul Remeika (1984, written communication) of the Vallecito basin and Impact Area was also incorporated into the compiled geologic map (Plate 4).

Topographic maps proved inadequate for accurate location in the badlands topography; 1:20,000 monochrome aerial photographs of the Soil

Conservation Service (1953) survey were therefore used in the field.

Aerial photo coverage of the study area by the SCS survey is summarized

in Appendix I. Stereo pairs were used for preliminary photogeologic

mapping and for field work. Plotted information was transferred to an

enlarged (1:12,000) topographic base using a stereo Zoom Transfer Scope

(Bausch and Lomb trademark). This method provides better registration

with topography than does the more conventional method of tracing from

a photomosaic, although it is more tie-consuming.

Measured sections

Detailed measured-section logs were prepared of some of the These better-exposed sequences in the FCV area (Figure 1.12, Plate 4). Boyles, graphic logs were prepared on a standardized form developed by in the field. In Scott, and Rine (1986); a scale of 1" = 5 m was used of mean grain these logs, the edge of the graphic column is a graph profile. Sedimentary size, rather than a depiction of the weathering and the nature of contacts are sketched structures, details of bedding, summarized in Figure in the field. Symbols used in these logs are 44 1.13. Sections were measured using the Jacob's staff method; a plane- of-sight Jacob's staff was developed during the study to facilitate oblique-to-dip measurement (Winker 1986).

Paleocurrent measurement

Paleocurrent directions (Appendix II) in sandstones were based

on foreset orientations in cross-stratification and ripple lamination,

and on bearings of linear features such a sole marks, flute casts,

parting lineations, ripple crests, and ripple troughs in plan view.

Single-rotation structural corrections for bedding dip were made by

stereonet in the field and later computed automatically; the tilt-

corrected paleocurrent directions are tabulated in Appendix II. In-

ference of true palocurrent direction also requires correction for

tectonic rotation about a vertical axis as revealed by paleomagnetic

declination anomalies (e. g. Johnson et al. 1983). Paleocurrent measurements were not made in conglomerates, although conglomerate data

from Kerr (1982, 1984) were incorporated in this study.

Thin-section petrography

Sandstone samples for petrography were collected from measured sections and from miscellaneous localities. Most of the samples were limited to the fine to lower-medium grained range (1.5 to 3 phi) to reduce the effect of grain size on compositional statistics (Ingersoll et al. 1984, Dickinson 1985). The samples ranged from friable, unce- mented sandstones to thoroughly carbonate-cemented and non-porous sandstones. 46

o sandstone section thickness Nr- in meters g vccmU v( 40 solid line: bioturbation estimated mean grain size (relative intensity) possible bioturbation • • dotted line: ...... maximum grain size (optional) ' C'-suite mixed provenance 35 dotted line: suites matrix size in ' L'-suite conglomerates and coquinas thin "I.. ' bed ..bedding-surface geometry reddish (may be foreshortened) slightly reddish thinly interbedded gray or claystone sandstone & mudstone olive-gray color 30 blue or dashed line: blue-green poorly exposed or G gray-green heavily weathered gravel (boulders to scale) not exposed ///// claystone intraclasts faults cutting section oyster shells

estimated gastropod shells stratigraphic equivalence miscellaneous shells NF-6-98 SEDIMENTARY oyster-shell coquina STRUCTURES r -s oyster shells 1,- in muddy matrix parallel lamination •.:=' 20 silicified wood 20 low-angle off= (with diameter in cm) cross-stratification plant fragments trough cross-stratification 7.71....n"-Za.. MC mud cracks tangential to tabular cross-stratification mud partings and large 'lasers convolute lamination calcareous nodules 5 small-scale ripple lamination CH E CM calcareous horizons climbing ripples (ripple drift) A 4.•4° Ey evaporite nodules Miser bedding sensu discrete burrows Reineck "I: wavy bedding Wunderlich BODY FOSS LS lenticular bedding 1968 - Os Ostrea sp. laminated mudstone - Oh Ostrea heermanni o pe pectinids rhythmite bedding ph pholads (poorly developed tu turritellids or exposed) -n ce ceritheids graded beds • Ba Balanus (RG-reverse grading) 4zo sd sand dollars' RO ABE -partial Bouma sequences TRACE FOSSILS AC 5 sole marks Ch Chondrites groove casts • Gy Gyrolithes flute casts Op Ophiomorpha Sc Scolicia flame structures Te Teichichnus ball-and-pillow structures Th Thalassinoides s cc- 'cone-in-cone' burrow debris-flow bed ('debrite") ct camel tracks 0

Figure 1.13. Explanation for measured-section logs in Chapters 4 and 5, and Plates 5 and 6. 46 All rock samples were vacuum-impregnated with translucent blue epoxy prior to thin-sectioning. One half of each thin section was etched and stained for potash feldspar (Bailey and Stevens 1960) and plagioclase (Laniz, Stevens, and Norman 1964). Staining of feldspars is deemed essential for unambiguous differentiation of untwinned feldspars from each other and from quartz, particularly when fine-grained sandstones are being studied.

Compositional statistics (Appendix III) were obtained from point counts using a counting stage; on each thin section a minimum of

400 framework grains was counted. Computation of framework-grain percentages was based on the Gazzi-Dickinson method (Ingersoll et al.

1984, Dickinson 1985), in order to further reduce grain-size effects on composition. In this method, individual quartz and feldspar crystals

>0.0625 mm (<4 phi) in apparent diameter within polycrystalline grains were counted as quartz or feldspar grains, rather than as lithic - grains.

Geographic nomenclature and numbering scheme

Topographic maps of the study area give relatively few geographic names. Additional names appear on published maps of the State

Park, and park rangers have coined a number of informal names; these were used when necessary. In a few instances it was necessary to coin new names. Geographic names used for the FCV section are shown in

Figures 1.3 and 1.4 and Plate 4.

Measured sections are designated by a two-letter geographic prefix (e.g. FC for Fish Creek Wash), followed by a number (e.g. FC-4); 47 samples or localities tied to measured sections are designated by a second number indicating the distance in meters from the base of the section (e.g. FC-4-105). Samples or localities not tied to a measured section are designated by a geographic prefix, a lower-case "m" (miscellaneous), and an arbitrary number (e. g. FC-m7-8). Locations of most measured sections are shown in Figure 1.12; all measured sections and

key miscellaneous localities are shown in Plate 4. CHAPTER 2

SANDSTONE COMPOSITION AND PROVENANCE

Field observations Field work in the FCV area revealed the presence of two

distinct sandstone types with local mixing. Recognition of these sand-

stone types is useful in stratigraphic subdivision and description, and

so will be discussed before stratigraphy. Merriam and Bandy (1965) had

previously recognized two discrete compositional suites in samples of

the FCV area and other localities in the Salton Trough, although their

samples were not precisely located with respect to geography or strati-

graphy. These suites were interpreted by Merriam and Bandy to represent

Colorado River and Peninsular Range provenances, which suggested the

testable hypothesis that the sandstone types Observed in the field

correspond to those documented by Merriam and Bandy (1965).

The first sandstone type recognized in the field is angular,

coarse to fine-grained, and ranges from poorly to well sorted. Bedding, lamination, and other sedimentary structures are typically well defined due to textural variability. The coarse-grained sandstones are commonly conglomeratic, and in some units grade laterally into conglomerates.

This type of sandstone is associated in some units with siltstone or mudstone. Under the hand lens, these sandstones appear angular and feldspar-rich. These sandstones are also typically biotite-rich,

48 49 especially in the finer grain sizes. Typical colors in fresh sandstones are gray, olive-gray, and greenish-gray; reddish sandstones are also common in the lower part of the section. Weathered surfaces may be more yellowish, brown, or rusty orange. The finer-grained, more biotite-rich sandstones are generally darker colored than coarser-grained sandstones, although color value is also affected by the amount of clay matrix. Associated siltstones and mudstones are typically dark olive- gray and fissile because of Abundant mica.

The second sandstone type ranges from very fine to fine (rarely

lower medium) in mean grain size, and from moderately well to very well

sorted. Under the hand lens, these sandstones are typically subrounded and appear to be more quartz-rich than the previous group. Pinkish or

orange grains are common, and mica is scarce or absent. In many outcrops, bedding, lamination, and other structures are poorly expressed or cryptic, apparently due to uniformity of texture and composition.

However, in some beds in the Palm Spring Formation, heavy-mineral concentrations accentuate parallel lamination, ripple lamination, and parting lineations. These sandstones are typically light yellowish gray

(Imperial Formation) to pale orange (Palm Spring Formation). They grade into coarse to fine siltstones of similar composition and color, and are typically associated with non-fissile gray to reddish claystones.

The siltstOnes and claystones, like the sandstones, have little mica.

In the rest of this paper the first sandstone type will be referred to as the "L" suite and the second as the "C" suite. In some contexts, these terms will also be used to include the 50 characteristically associated conglomerates, siltstones, and claystones.

In some beds, compositional gradations between the two sandstone types are observed. A common occurrence of mixed-composition sandstone is in oyster-shell coquina beds of the Imperial Formation; these coquinas are generally associated with "C"-suite sandstones but coarse sand, granules, and pebbles of the "L" suite may be concentrated along with shells and shell fragments. Another common occurrence is in upward-fining sandstone sequences in the Palm Spring Formation, in which coarser-grained, angular, gray sandstone ("L" suite or mixed "L" and "C") grades up into finer-grained, rounded, pale orange sandstone

("C" suite).

A third discrete sandstone type occurs in the basal unit at

Split Mountain (Red Rock member of this study). These sandstones are

typically very coarse-grained to granular and are richly feldspathic,

but are better rounded than typical "L"-suite sandstones. Large pink

feldspar grains, not Observed in the typical "L" suite, are common; suite. these sandstones are also less micaceous then the typical "L"

The appearance of these sandstones is evocative of reworked granitic Kerr grus. Mineralogic composition of these sandstones was studied by

(1982, 1984) and will be reviewed later.

Thin-section petrography "L"- Samples were collected for thin-section petrography from sandstones, based suite, "C"-suite, and a few intermediate-composition in which the on field identification. Sampled areas included sequences 51 two suites occur in stratigraphie alternation. Point-counting data and statistics are tabulated in Appendix III. Framework-grain compositions are summarized in standard QFL, QmFLt, and QmPK ternary plots (e. g. Ingersoll et al. 1984, Dickinson 1985) (Figure 2.1); other compositional data are summarized graphically in Figures 2.2 and 2.3. Lithic-

grain compositions (standard QmLsLv ternary diagrams) were not plotted because of the scarcity of lithic grains in "L"-suite sandstones and because of the difficulty of consistently differentiating lithic grain types in "C"-suite sandstones.

Framework-grain composition

Framework-grain compositions of "L"-suite and "C"-suite . sandstones plot in two discrete fields. The "L" suite consists of subequal

amounts of quartz and feldspar, with slightly more feldspar, and a P/K

ratio of -2:1 or more. "L" sandstones typically contain a few percent

polycrystalline quartz grains (some of them foliated), so that on a

QmFLt plot the field shifts slightly toward the Lt pole from its posi-

tion on the OFL plot (Figure 2.1). Plagioclase constitutes 62% to 85%

of the total feldspar content (Figure 2.2). The majority of plagioclase and K-feldspar grains are untwinned. Biotite, which by convention is not included with the framework grains, ranges from 11% to 32% of the framework-grain total (Figure 2.2). The biotite content is highly dependent on grain size, but is consistently greater than for the "C" suite.

The "L" suite plots within the arkose fields of McBride (1963) and Folk (1968), and is close to other first-cycle arkoses from 52

O 'C suite

ancestral • .1.." suite Colorado River source P mixed provenance

local basement source

Figure 2.1. Framework-grain composition of sandstones from FCV section, on standard OFL, QmFLt, and QmPK plots (e. g. Dickinson et al. 1985). 53

C/Q / • nnn Hp n 0 4 5 6 7 8%

P/F

ri " p c\icv 40 45 50 551 60, 65, 710 75 80 85% 10

M/S 73 J..... • • o 5 10 15 20 215 30 35%

CEMENT + POROSITY 111 • n Fr] Ill 6 lb 15 30 315 4:10%

mixed provenance *C' suite D 111 "L" suite

Figure 2.2. Miscellaneous compositional parameters of sandstones from FCV section. C/0 = chert/total quartz; P/F = plagioclase/total feldspar; M/S = mica/total sand (mica is mostly biotite).

Pd DC-7b-209 • 0 • LC-1-26 • • o • • Po LC-1-14 • • • • FC-4b-48 • • 0 • *. Pd FC-4a-10 • • 0 • • NF-4-47 • 0 • • CT-2-78 • • • Pch FC-3-17 • • 0 • • FC-2-110 • • o • • FC-1-26 • •0 • Pdg FC-7-110 • • o EK-2-3 • o • • NF-1-194 • loo • • FC-5-50 s • o • • Pw NF-1-19 • • o • EK-1-180 • 0 • • , 0 20 40 600 1 80%' 20 40%0 10' 20 ' 30 20 40% 40 0' 60 80% •L/L+Q P/F FIG Lv/(Lv+Ls) oLt/(Lt+Qm)

Figure 2.3. Miscellaneous compositional parameters for "C"-suite sandstones, plotted in stratigraphic sequence. 55 Southern California and to the "ideal arkose" of Dickinson (1985). This composition is characteristic of continental-block provenance resulting from basement uplift (Dickinson and Suczek 1979, Dickinson et al. 1983).

"C"-suite sandstones are less feldspathic and more lithic-rich

than the "L"-suite sandstones (Figure 2.1), with subequal amounts of

plagioclase and K-feldspar (Figure 2.2). As with fine-grained "L"-suite

sandstones, the majority of both plagioclase and K-feldspar grains are

untwinned. Many lithic fragments are difficult to identify, but a

number of unambiguously volcanic and sedimentary (chert and limestone)

lithic grains could be differentiated. In sandstones with no calcite

cement, carbonate clasts are absent. Mica content is <5% of the

framework-grain total in all examined thin sections of "C"-suite sand-

stones, and <1% in the majority of samples (Figure 2.2). Some

two-component statistics on "C"-suite sandstones were plotted against

stratigraphic position to test for systematic up-section variation

(Figure 2.3), but no systematic variations were observed.

Most "C"-suite sandstones plot within the "lithic arkose" field fields of of Folk (1968) and the "lithic arkose" to "lithic subarkose" provenance" McBride (1963). They also plot within the "recycled orogen

field of Dickinson and Suciek (1979) and Dickinson et al. (1983), and suites" of Dickinson are similar in composition to "collision belt sand

(1985). 56

Other characteristics

As Observed in the field, "C"-suite sandstones are consistently better rounded than "L"-suite sandstones; this difference is readily apparent in thin sections. Feldspars in "C"-suite sandstones are generally somewhat rounded, unlike the angular feldspars of the "L" suite, suggesting that the "C" suite composition is not a result of mixing

"L"-suite sands with a more quartzose and lithic-rich end member. An

occurrence of both coarse, angular feldspars and fine, rounded feld-

spars in the same thin section is a good indicator of mixed "L" and "C"

suites.

Four of 16 "C"-suite thin sections contained calcite-filled

tests of foraminifers. As with carbonate-lithic clasts, foraminifers

only occur where calcite cement is present. Merriam and Bandy (1965) foram- noted the presence of seven species of similar calcite-filled

iniferal tests in their sand samples from the Salton Trough. They the Mancos identified these as late Cretaceous species, and suggested presence Shale on the Colorado Plateau as the most likely source. The as a diagnostic indicator of these foraminifers has come to be regarded

of Colorado sands in the Salton Trough (e.g. Lucchitta, 1972). However, subject to dissolu- because such foraminifers are relatively scarce and source. tion, their absence does not rule out a Colorado River the presence Another characteristic of "C"-suite sandstones is on quartz grains; such of rare but distinctive rounded overgrowths

rounded overgrowths typically indicate a recycled sedimentary 67 provenance. Presumably, more recycled overgrowths could be identified by cathodoluminescence.

Examination of "C"-suite sandstones and disaggregated sands indicated that many of the quartz grains have a light reddish, brown, or orange coating, presumably hematite. These coatings probably account for the pink and orange grains observed under the hand lens and for the typical pale orange color of the sandstones in outcrop. These coatings may reflect recycling of sedimentary redbeds.

Diagenesis

Although diagenesis was not a major focus of the present study, sediment alteration is of interest because of the high degree of alteration observed in samples from geothermal wells in the Salton Trough

(Muffler and White 1969, Elders 1979b, McDowell and Elders 1979, Vonder

Haar and Howard 1981). The most obvious diagenetic features Observed in

this study are physical compaction by deformation of labile grains, carbonate cementation, hematite cementation, feldspar dissolution, and

carbonate dissolution. Recycled diagenetic features previously men-

tioned, specifically quartz overgrowths and hematite coatings, are

observed in the "C"-suite sandstones.

Physical compaction is most apparent in "L"-suite sandstones,

and has occurred primarily through ductile deformation of biotite grains. Hematite cementation is also apparently limited to "L"-suite sandstones. It is manifested primarily by the reddish coloration of some sandstones and sand matrix of conglomerates, particularly in units exposed in Split Mountain. Walker (1967) studied similar Neogene 58 arid-climate redbeds in the southwestern U. S. and demonstrated that hematite cement is derived mainly from chemical degradation of mafic minerals such as biotite.

Some carbonate content is virtually ubiquitous in Neogene sedimentary rocks of the FCV area, particularly in the finer-grained sediments, although carbonates may be absent from weathered rock. Car- bonate cements occur in both "C"-suite and "L"-suite sandstones.

Carbonate cements are mostly sparry and locally poikilotopic. In some

cases carbonate cement has replaced parts of feldspar grains; "holes"

in the grain framework suggest that some feldspar grains may have been

entirely replaced. Carbonate lithic grains are recognized primarily by the contrast between the fine-grained calcite of the clasts and the coarse-grained calcite of the cement. Sandstones that lack carbonate cement also lack carbonate lithic clasts, suggesting that absence of carbonate is due to dissolution, possibly on the outcrop.

Although most feldspar grains are sufficiently intact to permit unambiguous identification when stained, and the majority of feldspar grains appear relatively fresh and unaltered, a wide variety of alteration states was observed. These range from sericitization, to partial carbonate replacement, to extensive corrosion or dissolution as evidenced by wisps and skeletons of feldspar grains.

Sandstone porosity varies over a considerable range, from essentially zero to nearly 30% (Figure 2.4). Porosity is lower on the average in "L"-suite sandstones than in "C"-suite sandstones. Porosity also tends to decrease with depth, but this may reflect the fact that 59

"L"-suite samples are concentrated lower in the section. When total porosity plus carbonate cement is plotted (Figure 2.4), variations with composition and depth are much less and may not be significant. There- fore, porosity variations are caused largely by variability in the intensity and/or preservation of carbonate cements.

Comparison with other studies

Direct statistical comparison with results of other studies is hindered somewhat by lack of standardized techniques and procedures,

different philosophies on point-counting and statistical computation,

and inadequate documentation. Nevertheless, significant similarities

between the results of this and other studies can be found. In some

cases, plutonic lithic fragments reported in previous studies were

recalculated as quartz and feldspar using assumed ratios. This was done

to simulate the Dickinson-Gazzi method and thereby reduce grain-size

effects on compositional statistics and obtain more directly comparable

data. A ratio of 40% quartz to 60% feldspar was assumed in all such

recalculations, based on typical compositions for first-cycle arkoses

from Southern California (Dickinson 1985). of modern Van Andel (1964) made the first comprehensive study the first sand composition in the Gulf of California, and provided Although his detailed di.scussion of Colorado River sand composition. data on framework emphasis was on heavy minerals, he also presented the 0.062-0.25 mm grain composition. Ris point counts were limited to to the present study (2-4 phi) range, so grain-size effects relative up to 11% and should be minor. Plutonic-lithic contents 60

DC-7b-209 FG-m-8'L' LC-1-26 LC-1-14 FC-4b-48 FC-4a-10 FC-3-7 CT-2-78 CT-2-26 IL CT-la-14.5 IL FC-2-110 IM(C) FC-1-26 NF-4-47 EK-2-3 FC-7-110 NF-1-194 EK-1-166'C SM-1-140 IL NF-1-92(N) IL FC-5-50 NF-1-19 IC EK-1-18 IC EK-1-3.5 SM-3-57 IL SM-2a-7.5 SM-2-54 IL SM-2-21 IL CC-m-13 IL SM-mr5 IL porosity I 1 cement

Figure 2.4. Sandstone porosity and cement in FCV section, plotted in approximate stratigraphic sequence. 61 metamorphic-lithic contents up to 13% were reported. These were not recalculated; if they were it would shift the plotted positions away from the L pole on the QFL ternary diagram (Figure 2.5). Replotting of

Van Andel's (1964) data (Figure 2.5) shows three end-member composi-

tions: an "arkosic" end-member (similar to the "L" suite of this study), a "graywacke" end member, dominated by volcanic lithics, and a

"feldspathic" end member (similar to the "C" suite of the present

study). Many data points plot between the "arkose" and "graywacke" end members; these correspond to the "magmatic arc provenance" field of

Dickinson and Suczek (1979) and Dickinson et al. (1983). Van Andel

(1964) showed that distribution of the "arkose" and "graywacke" suites

in the Gulf is closely related to the distribution of Mesozoic plutonic

and late Cenozoic volcanic outcrops, respectively, in drainage basins bordering the Gulf. The "feldspathic" end member composition occurs

only in the northern part of the Gulf near the Colorado delta: within

the Colorado estuary, on the adjacent continental shelf, or on the

continental slope basinward from the delta. On closer examination, Van

Anders (1964) "arkosic" and "feldspathic" end members are not identi-

cal in composition to the "L" and "C" suites of the present study; his data plot closer to the 43 pole on a QFL diagram (Figure 2.5). Whether

this discrepancy is due to actual compositional differences, to differences in technique, or to misidentification of unstained and untwinned feldspar as quartz, is not apparent from Van Andel's (1964) discussion. 62

modern Colorado suite, Salton Trough (Van de Kamp, Colorado suite, 1973) Gulf of California (Van Andel, 1964)

local basement, Salton Trough mixed local (Van de Kamp, N volcanic & plutonic 1973) 50 (Van de Kamp, 50 1973)

50 Red Rock member at Split Mountain °rep,

1Red Rock member L_ Elephant Trees member at Fish Creek Mts. at Split Mountain v , / 50

Figure 2.5. QFL compositional fields from this study (Figure 2.1) compared with those of previous studies. Data from Kerr (1982) are shown before and after recalculation of plutonic lithic grains as 60% feldspar and 40% quartz; stratigraphic nomenclature is from this study. 63

Merriam and Bandy (1965) first convincingly demonstrated the presence of Colorado River-derived sands in Neogene strata of the

Salton Trough; they also presented data from modern sediments. They recognized two discrete compositional fields, attributed to Colorado

River and Peninsular Range (i. e. more local) sources (Figure 2.6).

However, these two fields plot significantly differently from the "C"

and "L" fields of the present study, and from the "feldspathic" and

"arkose" end members of Van Andel (1964). Again, it is not clear

whether these compositional differences among studies are real or an

artifact. Merriam and Bandy's (1965) data show considerable overlap of

compositional fields between modern Colorado River sands and Neogene

Colorado-derived sandstones (Figure 2.6); these compositions are essen-

tially indistinguishable.

Merriam and Bandy's (1965) data suggest the following charac-

teristics which could be used to distinguish a Colorado River source

from Peninsular Range sources:

1. Framework-grain composition;

2. Presence of Upper Cretaceous foraminifers with calcite--

filled tests reworked from the Mancos Shale; other sedimen- 3. Presence of calcite or dolomite fragments or

tary lithic fragments; and plagioclase, 4. Scarcity of biotite, hornblende, augite, Ranges but rare in the which are common in rocks of the Peninsular a higher Colorado delta (results of the present study indicated 64

modern Colorado (Merriam & Bandy, 1965)

Pliocene Colorado (Merriam & Bandy, 1965)

Quaternary Colorado (Potter, 1978)

local basement (Merriam & Bandy, 1965) 50 50

modern Colorado (Van de Kamp, 1973) Quaternary Colorado (Potter, 1978) Pliocene Colorado (this study) modern Colorado (Merriam & local plutonic Bandy, 1965) (Van de Kamp, 1973) Pliocene Colorado (Merriam & local basement 50 Bandy, 1965) (Merriam & 50 Bandy, 1965) mixed local plutonic & volcanic local plutonic (Van de Kamp, (this study) 1973)

Figure 2.6. QmFLt and QmPK compositional fields from this study (Figure 2.1) compared with those of previous studies. 65 percentage of plagioclase in the "C" suite than recognized by Merriam and Bandy).

Muffler and Doe (1968) studied Quaternary and Neogene sediment in the Salton Trough from a geochemical and isotopic standpoint. Their data corroborate Merriam and Bandy's (1965) evidence for a Colorado River source. Lead isotopic data indicate a mean age for Salton Trough detritus of 1700 Ma (Proterozoic), which would require a source from

either very old basement or from polycyclic sediment. This apparent

isotopic age cannot be accounted for by the relative young basement rocks surrounding the Salton Trough.

Van de Kamp (1973) presented petrographic data from modern

sands of the Salton Trough. His data included high percentages of

plutonic-lithic fragments, and were therefore recalculated for plotting

here (Figures 2.5, 2.6). Van de Kamp (1973) recognized compositional

fields similar to those of Van Andel (1964) for the Gulf of California.

Specifically, he recognized a continuous range between a "local base-

ment" end member (similar to Van Andel's "arkosic" sands and the "L"

suite of the present study) and a "local volcanic basement" end member

(similar to van Andel's "graywacke" sands). Van de Kamp (1973) also

recognized a discrete "modern Colorado" field, similar to the "C" suite

of this study. Again, however, end-member compositions vary among these

studies.

Potter (1978, also written communication to W. R. Dickinson,

1979) presented compositional data for sands from a number of Quater- nary rivers including the Colorado. His Colorado River sample was taken 66 from 26 m below the surface in Shell Development Corp. core R4419,

Section 22, T 10 S, R 25 W, Yuma County, Arizona, which is located along the lower Colorado valley above Yuma. Potter's samples were limited to fine to medium sands where possible, and point counts were made on impregnated thin sections stained for two feldspars, as in the present study. His data point for the Colorado River (61Qm:21F:18Lt;

74Qm:9P:17K) plots on QmFLt and QmPK diagrams (Figure 2.6) within the

"C"-suite field of the present study, but not within the "Colorado"

fields of Merriam and Bandy (1965) or Van de Kamp (1973).

Other petrographic data from Salton Trough sands and sandstones

have been reported by Bird (1975) and Wagoner (1977). Bird (1975)

reported on the DWR #1 borehole in the Dunes hydrothermal system, which

is located on the Colorado delta plain. All compositions of these

presumably Quaternary sands plot near the "Colorado River sands" fields

of Van de Kamp (1973) and Merriam and Bandy (1965). Bird's composi- suite of the tional data plot farther from the F pole than does the "C" Pleistocene present study. Wagoner (1977) presented data from the in the (?) Brawley and Borrego Formations in the San Felipe Hills cover a large western side of the Salton Trough. His compositional data the arkose- area on a QmFLt diagram, but are all more quartz-rich than other studies. litharenite spectrum of locally-derived sands of River Therefore, his data are consistent with a dominant Colorado

source. sandstones from the Kerr (1982) reported petrographic data on from Split Mountain FCV section, including his "braided-stream facies" 67 and Red Rock Canyon in the CIA (Red Rock member of this study), and his

"alluvial fan facies" in Split Mountain (Elephant Trees member, lower fanglomerate unit of this study). Compositional data were presented only as averages and ranges rather than as counts on individual thin sections. Sandstones studied by Kerr (1982) were coarse-grained and dominated by plutonic lithic fragments; to Obtain statistical data comparable to those of the present study, plutonic lithic clasts were recalculated as 40% quartz and 60% feldspar. The recalculated data from both stratigraphic units plot on a QFL diagram (Figure 2.5) in the arkose field of both McBride (1963) and Folk (1968), and between the F pole and the "L"-suite field of the present study. Feldspar composition

of Kerr's (1982) "alluvial fan facies" is characterized by a dominance

of plagioclase over K-feldspar (average ratio 80:20), similar to the

"L" suite of the present study. His "braided-stream facies" sandstones

have subequal amounts of plagioclase and K-feldspar, which differen-

tiate these sandstones from the typical "L" suite.

Provenance suites of Four end-member compositions characterize Neogene sandstones

the Salton Trough and Gulf of California: but 1. A Colorado River composition, relatively quartz-rich, sedimentary mixed with plagioclase and K-feldspar and volcanic and This suite lithic grains. This is the "C"-suite of the present study.

can also be recognized on the basis of other criteria, including reworked calcite-filled late rounding and sorting, the presence of (Figure 2.7). Cretaceous foraminifers, and recycled quartz overgrowths 68

THIN FIELD CUTTINGS SECTION 1. FRAMEWORK-GRAIN COMPOSITION •

2. ROUNDING & SORTING • •

3. QUARTZ OVERGROWTHS (RECYCLED) •

4. REWORKED CRETACEOUS FORAMS •

5.BIOTITE CONTENT • •

6. PINK & ORANGE SAND GRAINS

Figure 2.7. Criteria for recognition of Colorado River-derived sand in Salton Trough and Gulf of California. 69 2. A plagioclase-rich arkosic end-member. This is the "L" suite of the present study. The source of these sands is readily attributed to granodiorites, quartz diorites, and tonalites of the Peninsular Ranges, including local mountains such as the Vallecitos. 3. A litharenite end-member dominated by volcanic-lithic clasts. The dominant source of these sands is probably Cenozoic volcanic rocks of the Salton Trough and circum-Gulf regions. Many data points plot along a mixing line between this end member and the plagioclase-rich arkose end member. Sandstones of the volcanic- litharenite end member and litharenite-arkose mixing-line apparently do not occur in the FCV section, but do occur in Neogene sandstones in the

Coyote Mountains (Carrie Campen-Bell, 1985, oral communication), and are included in this paper in the locally-derived ("L" suite) provenance suite.

4. An arkosic end member with subequal plagioclase and

K-feldspar. This composition occurs in the basal unit of the Split

Mountain-FCV section (Red Rock member of this study), but it may be more extensive. A source area has not been identified, but some associated gravel clasts suggest an extra-local source, and paleocurrent directions suggest derivation from the south or east (Kerr, 1982).

Recognition of subtler compositional differences is hampered by nonuniforMity of procedures and documentation, and possibly by errors in grain identification. These problems are reflected in discrepancies among various studies in compositional statistics for what are presumably the same provenance suites (Figures 2.5, 2.6). Some of these 70 discrepancies may be due to random counting error (Figure 2.8), but this explanation alone is not convincing. A more likely source of systematic error is misidentification of unstained and untwinned fine-

grained feldspar as quartz. Of the previous studies reviewed here, only

Potter (1978) mentioned the use of staining for two feldspars. His

results are consistent with those of the present study, whereas other

studies show significantly less feldspar in the Colorado River suite.In the present study, plagioclase was found to more abundant in "C"-suite

sandstones and K-feldspar more abundant in "C"-suite sandstones than reported by Merriam and Bandy (1965). Systematic differences in

reported percentages of lithic fragments could result from grain-size differences, if the Dickinson-Gazzi method is not used (e.g. Kerr

1982). With standardized techniques it might be possible to discern subtle differences in local source areas, mixing of different provenances, and changes through geologic time of sediment carried by the Colorado River.

In spite of such problems, results of these studies clearly show the presence of distinctive Colorado River-derived sand in late

Cenozoic sediments of the Salton Trough and northern Gulf of Califor- nia; the present study indicates that this distinction can be readily made in the field. Therefore, the stratigraphic section of FCV and adjacent areas can be subdivided on the basis of provenance (Figures

2.9, 2.10), which provides an alternative to the conventional subdivision into marine and nonmarine units. 71

Q/Qm

I statistica4l 0c0ountiingtserror

(after Van der Plas & Tobi, 1965) /OKA\ %%\,. %%\,. 5v o )11- (0>çc). PXC))(CDX(DA Ç) <(=.2<0) ,ÇD)(c)) ç=),<<=> çc:) n n 1-4 1-4 F—Y 6-1 I I 1--1n1-1 4—fnol-4 I I I L/ 50 Lt

directions of shift direction of shift due to use of due to misidentification Dickinson-Gazzi method of feldspar as quartz (dashed - average arkose)

Figure 2.8. Possible sources of discrepancies in sandstone compositions between this and previous studies. Statistical counting error refers only to variance due to finite sample size, not identification errors. 72

alternating or interfingering 'C' and 'L'

arkose (P=10 with some extra—local clasts

(Mt Is authigenic)

Figure 2.9. Geologic map of FCV section (Figure 1.5), showing distribution of provenance suites. 73 CHAPTER 3

STRATIGRAPHY OF THE LOWER NONMARINE UNITS

Stratigraphy of the FCV section and adjacent areas has

traditionally been subdivided into three broad categories (Figures

1.5-1.11): a lower sequence of predominantly nonmarine units (Split

Mountain, Anza, and Alverson Formations); a middle sequence of marine

units (Imperial Formation and part of Split Mountain Formation); and an

upper sequence of nonmarine units (Palm Spring, Borrego, and Brawley

Formations, and Canebrake and Ocotillo Conglomerates). Stratigraphic

descriptions in this paper are organized in the same three categories.

Tarbet and Holman's (1944) original definition of the Split

Mountain Formation included marine facies (Lycium and Stone Wash mem-

bers of the present study) as well as nonmarine facies. Most subsequent

workers followed this usage (Tarbet 1951, ver Planck 1962, Dibblee

1954, Smith 1962, Woodard 1963, 1974), although Stump (1972) reassigned

those marine facies to the Imperial Formation; this paper will follow

Stump's lead (Figure 1.9). The Fish Creek Gypsum could arguably be

included under either marine or nonmarine stratigraphy; it will be

discussed here with the nonmarine units. Field work in this study

concentrated on the marine and upper nonmarine units. Therefore,

description of the lower nonmarine units is based largely on previous

work (Ruisaard 1979, Mace 1981, Gjerde 1982, Kerr 1982, 1984, Kerr et

74 75 al. 1979), supplemented by observations from the author's reconnaissance.

Existing stratigraphic nomenclature was found to be inadequate and even misleading for identifying and describing genetically significant units. In particular, uses of "Split Mountain" are inconsistent

among different authors and different localities; to attempt redefinition of this unit would probably just add to the confusion. Therefore, new, informal names, Red Rock member and Elephant Trees member are used

in this paper for nonmarine strata formerly assigned to the Split

Mountain Formation. Stratigraphic relationships among the lower nonmarine units are summarized in Figures 1.6 and 3.1.

Red Rock member (Mk) (new name)

The Red Rock member consists of white, pink, orange, and red

arkosic sandstones, mostly coarse-grained and conglomeratic, which

occur in discrete lenses between underlying plutonic and metamorphic

basement and overlying fanglomerates (Elephant Trees member) and vol-

canic rocks (Alverson Formation). This unit has previously been

included in the Split Mountain Formation (Tarbet and Holman 1944) or

Anza Formation (Woodard 1963, 1974). It is regarded here as genetically

distinct from the overlying fanglomerates and important in regional

history, and so has been given a separate designation. It occurs in

three main areas, separated by pinchouts (Figure 1.6).

1. Split Mountain. The lowest sedimentary unit in Split

Mountain is a very coarse to granular, thick-bedded to massive, conglomeratic sandstones of uniform pinkish brawn (2-5 YR 5/6) which 76

ou 0 ..t U 05 4-1 E M .44 01 c 1.) 4 OE 05 •

§ cd" 0 +) •,-1 3 n•-•I o a —1 c1-1 o 05 0 00 •ri • 4.) g U rn

c. o 05 (t 0 U

L.4 05 tl) g

01 '00 0 5 a a) o Cli 01 E •,-1 1-4 •,-1 U X

0 0 Q. 0 0 CI

• .4.> G 0 0 c.; 0 X (1) S., 4 t) 05 • 1•.( ;LI C.) •ri X 4-) U ••-I (1) Cm+ 0 77 weathers to a lighter pink (5 YR 8/2-4). The texture is very distinc- and is tive suggestive of a reworked and rounded granitic grus. Previous studies of Split Mountain have included this unit in the "red member" of the Split Mountain Formation (ver Planck 1952, Smith 1962), or have designated it the "lower member" (Woodard 1963, 1974) or "braided-stream facies" (Kerr 1982, 1984) of the Anza Formation; it is reassigned here to the Red Rock member.

Maximum thickness is -185 in just east of Split Mountain Gorge,

in the axis of the Split Mountain Anticline (Kerr 1982). In the Gorge

itself, only the uppermost part of the Red Rock member is exposed; west

of the Gorge this member is absent and Elephant Trees fanglomerate rests directly on basement. Approximately 16 in of Red Rock member is exposed in Cairn Canyon, but the member is absent at Boundary Mountain; the intervening pinch-out is concealed by colluvium. The Red Rock member is also absent in the vicinity of the U. S. Gypsum quarry, where conglomerate and gypsum directly overlie basement. The contact with overlying fanglomerate is generally gradational over a few meters, but is fairly sharp and well-defined relative to the total thickness of the units. An angular discordance of -10-15° exposed in Split Mountain

Gorge is actually a few meters below the lithologic change to conglomerate; it is regarded as a local intraformational unconformity.

2. Carrizo Impact Area. Sandstones below the Alverson volcanics in the CIA display a greater variety of coloration and lithofacies than the basal sedimentary unit in Split Mountain, but are similar in most respects and are therefore assigned to the same unit. Previous workers 78 have assigned these sandstones to the Split Mountain Formation (Smith 1962, Ruisaard 1979). They are typically white, gray, or pale pink, but are locally pink to red-orange, particularly below the contact with Alverson basalt. The colors are particularly vivid in Red Rock Canyon and Pink Valley. In addition to the dominant lithofacies of very

coarse, conglomeratic sandstone, the Red Rock member in the CIA also

includes a unit of basalt-clast conglomerate, and a unit of medium-

grained sandstone with very large-scale cross-stratification, both of which are exposed in Red Rock Canyon (Kerr 1982).

The Red Rock member in the CIA is generally -180 m thick

(Ruisaard 1979, Kerr 1982), but it pinches out northwest of Red Rock

Canyon and east of Barrett Canyon. North of Red Rock Canyon, basalt-

flow tongues interfinger with this member, whereas south of Red Rock

Canyon the Red Rock-Alverson contact is an angular discordance of 10-13° (Ruisaard 1979).

3. Coyote Mountains. Coarse, conglomeratic, arkosic sandstones with lenticular conglomerate beds occur in two areas in the Coyote

Mountains, above metamorphic basement and below basalt of the Alverson

Formation (Christensen 1957, Ruisaard 1979). As in the CIA, this unit has previously been designated the Split Mountain Formation. Since this unit has the same stratigraphic position as and similar lithology to the Red Rock member in the Carrizo Impact area, the same name is used here. In Fossil Canyon, the Red Rock member is up to 31 n thick and is separated from the overlying Alverson basalt by an angular discordance up to 16 , whereas in Ocotillo Canyon the Red Rock member is >140 m 79 thick and interfingers with basal Alverson basalt flows (Ruisaard

1979). The Red Rock member is absent in Painted Gorge and vicinity, where Alverson volcanics rest directly on metamorphic basement, and in the central Coyote Mountains, where the Imperial Formation rests on basement. The Red Rock member and Alverson Formation also crop out locally in the western Coyote Mountains, east of Canyon Sin Nombre.

Sedimentary facies

Kerr (1982, 1984) studied the sedimentology of the basal

nonmarine-clastic section of Split Mountain and the Carrizo Impact

Area, including the Red Rock member of the present study (his "braided

stream facies"). According to Kerr (1982), this unit in Split Mountain

consists of lenticular, upward-thinning cycles up to 6 m in thick-

ness. The typical cycle reportedly consists of (1) basal,

clast-supported cobble conglomerate in lensoid beds; (2) structureless

or crudely parallel-laminated granular, very coarse-grained sandstone

in beds up to 1.5 m thick, or very large-scale, cross-stratified

sandstone; and (3) sequences of trough cross-stratified, tabular cross-

stratified, and structureless or crudely parallel-laminated

sandstones. overstated the In the author's opinion, Kerr (1982, 1984) member in Split degree of organization in bedding of the Red Rock

Mountain. Cobbles occur in more or less discrete, lenticular con- upward-fining glomerate beds which could be construed as the bases of in varying or upward-thinning cycles; cobbles also occur dispersed structures are concentrations throughout the sandstones. Sedimentary 80 generally not well developed or expressed in outcrop, and the sandstones mostly appear to be very thickly to massively bedded. Overall, the Red Rock member appeared to be fairly disorganized, with a somewhat random distribution of the bedding types described by Kerr (1982,

1984).

In the CIA, Kerr (1982, 1984) reported thinner, lenticular,

cyclic units averaging -1 in thick. In these cycles, the typical cycle

reportedly consists of (1) tabular cross-stratified sandstone; (2)

structureless or crudely parallel-laminated sandstone; and (3) trough

cross-stratification (or a mixture of all three bedding types). The

uppermost part of the Red Rock member consists of well sorted, medium-

grained sandstone, characterized by large-scale wedge

cross-stratification in sets up to 4 in thick. Sand grains in this

facies are generally frosted and subrounded.

Ruisaard (1979) described sections of the Red Rock member (his

Split Mountain Formation) in the Red Rock Canyon area and in the Coyote

Mountains. North of Red Rock Canyon, he reported "lahar deposits" in an which he characterized as "subrounded-to-angular volcanic clasts the clast population immature ... sandy matrix. Volcanic rocks dominate

with minor amounts of basement rock." Near the top of the Red Rock with member, basalt-flow tongues of the Alverson Formation interfinger in beds up to Red Rock sandstones. These sandstones are coarse-grained,

1 in thick, with volcanic pebbles and cobbles, alternating with thinner

(up to 20 cm ) lenses of medium to fine-grained sandstone. Above that 81 sequence is 82 m of "plutonic clast megabreccia" which will be discussed under the Elephant Trees member.

Provenance and paleocurrents

The Red Rock member is unique among units of the FCV-CIA-Coyote

Mountain area in terms of paleocurrent directions (Figure 3.2), sand-

stone composition (Chapter 2), and clast composition. Gravel-size

clasts in the Red Rock member are derived from three distinct

sources: (1) local plutonic and metamorphic basement, as indicated by

granitoid igneous, gneiss, marble, and amphibolite clasts; (2) Miocene

basalt of the Alverson Formation or similar units; and (3) an extra-

local source indicated by metavolcanic and metasedimentary clasts

similar to those of the Eocene Ballena Gravels (Kerr 1982, 1982, Kerr

et al. 1979). Dominant paleocurrent directions in Kerr's "braided-

stream facies" are to the north in Split Mountain and to the northwest

in the Carrizo Impact Area (Figure 3.2). of Near Split Mountain Gorge, the Red Rock member consists

95-87% clasts of quartz diorite, tonalite, and pegmatite, and 5-10%

metamorphic clasts, including metavolcanic clasts similar to those of of the Red Rock the Ballena Gravel (Kerr 1982). In the upper third

member, basalt clasts also occur, probably derived from the Alverson composition of Formation (Kerr 1982). Kerr reported an average QFL

31:19:50%; the recalculated composition (Chapter 2) is 50:47.6:2.5%,

which is within the "L" field of the present study. Feldspar composi-

tion is anomalously potash-rich, with an average P/F ratio of 50%. In granules of pink Split Mountain, the Red Rock member contains abundant 82

'braided-stream facies" 'braided-stream facies" Carrizo Impact Area Split Mountain

1111 foresets - cross- stratification gravel a-axis 20% bearing

channel axes

'eolian facies" Carrizo Impact Area

Figure 3.2. Paleocurrent roses for Red Rock member (nomenclature of present study) from Kerr (1982), using his paleoenvironmental interpretations. 83 feldspar, probably K-feldspar. Lesser components of metamorphic lithic (0-6%; average 2%) and volcanic lithic (0-2*, average 0.5%) clasts were also reported by Kerr (1982). Foreset orientations in cross-stratification (N=36) in Split Mountain indicate a strong northerly dominance (Figure 3.2).

In the CIA, gravel clasts of Kerr's (1982) "braided-stream facies" consist of subequal amounts of plutonic (49%) and basalt (42%)

cleats, with 9% of quartzite, hornfels, and schist clasts. The latter

may be derived from local metamorphic basement. Average QFL composition

of the sandstone is 18:45:37, with exclusively plutonic lithic frag-

ments; the recalculated composition is 33:67:0, or slightly outside the

"L"-suite field of the present study. The average P/F ratio is 44%,

well outside the "L" suite field of the present study but similar to

the Red Rock member in Split Mountain. Foreset orientations show a

broader scatter than in the Split Mountain area, but the mean direction

is approximately to the northwest (Figure 3.2). Foreset orientations of

the "eolian facies" range over 1500, but are dominantly to the south to

southeast (Figure 3.2).

Kerr's data do not address the problem of vertical variation in

provenance of the Red Rock member in the CIA, particularly with regard

to Alverson-derived cleats. Ruisaard (1979) reported "lahar deposits"

with predominantly volcanic cleats within the unit north of Red Rock

Canyon but the precise stratigraphic position was not given.

South of Red Rock Canyon, he reported "a few volcanic cleats locally" 84 within the unit, and at Fossil Canyon, "volcanic clasts ... intermixed

within the upper section."

Lithologically equivalent units

The Table Mountain Gravels, originally defined by Miller

(1935), occur in the Jacumba area near the Mexican border (Plate

3). They disconformably overlie a granitic-metamorphic basement similar

to that in the PCV area, and are overlain by the lower Miocene Jacumba

Basalts, which are probably equivalent to the Alverson Formation

(discussed in the next section). Minch and Abbott (1973) described the

Table Mountain as light brown, friable, medium to coarse-grained sand-

stone and conglomeratic sandstone. Clast composition was described as

-50% low-grade metavolcanic and metasedimentary (evidently similar to

the extra-local clasts described by Kerr 1982), -25% sandstone and

quartzite, and -12% local granitic. The contact with the overlying

Jacumba Basalt was described as a highly eroded, irregular basal con-

tact. On the basis of stratigraphic position, lithology, and the

presence of extra-local clasts, the Table Mountain Gravels may be

similar or equivalent to the Red Rock member.

Alverson Formation (Ma) basalt flows, with The Alverson Formation consists primarily of

associated andesitic to dacitic tuffs, agglomerates and volcaniclas- the Split Mountain tics, and lava domes and plugs. It is absent from in the area and the central Coyote Mountains, but is well exposed Gorge to west of eastern Coyote Mountains from the mouth of Painted 85

Fossil Canyon, locally in the western Coyote Mountains east of Canyon

Sin Nombre, and in the CIA from Boundary Mountain to Barrett Canyon

(Figures 1.6, 3.1, Plate 4). Tarbet and Holman (1944) named this unit the Alverson Canyon Formation (the type locality is now more commonly known as Fossil Canyon); Dibblee (1964) substituted the somewhat misleading name Alverson Andesite. Stratigraphy of the Alverson volcanics

(Figures 3.3-3.6) has been described in considerable detail by Ruisaard

(1979); the following discussion is based largely on his work, with

supplemental data from Mace (1981) and Gjerde (1982).

Over much of its outcrop area the Alverson overlies the Red

Rock member, typically with angular discordance of as much as

16°. Locally, as in Painted Gorge, the Alverson rests directly on

metamorphic basement. The Alverson is overlain in most areas by the

Latrania member of the Imperial Formation, with contacts ranging from

gradational to angular discordance up to 200. Between Red Rock Canyon out and Boundary Mountain, the Alverson interfingers with and pinches

into the upper part of the Red Rock member, and is overlain by debris-

flow beds and a megabreccia of the Elephant Trees member (Figure lava flow; the 3.6). The base of the Alverson is defined by the first as a flow, lapilli top is defined by the uppermost volcanic bed, such

tuff, or crumble breccia. short Total thickness of the Alverson varies greatly over measured thickness is -140 m distances (Christensen 1957). The maximum reported by near Red Rock Canyon (Figure 3.6); a thickness of 279 m on a composite of Ruisaard (1979) for the Fossil Canyon area is based 86

Figure 3.3. Index map of four outcrop areas of Alverson Formation, in vicinity of Fossil Canyon, Coyote Mountains, from Ruisaard (1979). Location of canyons in Coyote Mountains is shown in Figure 4.2. • 3

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I .1:14 b°{1) 4-' .00 I.., IV ° I -0 .,..4= Ci VI I 2 i 4., o ..-I• cr a) -0 ...... E I ..v ,..-i 1-. 2a C..) g œ CI) I I I I 0:1 a) CT) 2 tu ogI I I i.., .4' 5 I aZt 44 ".1 ^C11) 5 .1-14.' 0 20 E I 7:$ c.., c.) 4) > 0 s., .c 0 = X x .> co . 0 .ro a E = = g -a-, d •O,-1 0 0.r4 —C9 ... To ii 7 0 c.1 ô I.:, :2 E r!e*".. ci 1 al om,z cc F2 .. cc E I go• -'-- r2 '.I0 c. I 41) '-' 4 C.) 0 = 4,) Vs., L....CI.> .1- 0 2 .() b .., '..-014 ,...... , t. 4 Ptl) 09, 0 P -, "C) r-I C.., I \ / / -4') -.4 1,0 5., ^0 clea / u2 / ( 1 ) g 4 ..4 / / -0 4 ) , -4 + I ) 2 / .5 / 9) at / 4-' 1:, s... g C4 / 4-4 a) ou / / O ta in 0 / 0 in 02 \ . o 1-, œ / ..,7, \ / 00 g 11. / . / gm 4 ""Ic,,, \ 7 ? c.; 6 4-1 0.8 ,-;.. \ / 5 - M a, cc, 2 2 2 0 5 \ 1 g / a) --44-, Oco I 2 g O -Hu) "Ti. g ""' DO 2 v / "-4 -4. O G) 4,r-4 : 'DO: .,C.21 gn,„ 0 co) Q

03 r'34 .op.I s., ..._....'") .... 0.) 4-) 'DO 4 1,1 eri 4,-) E ' 1..1 •4 I I 1-1 01 o 8 I 0 0 I fc1.., u 5-1 ""o -0 0 .0 t) .0 ti) ".4 Cr E 0 — .. .4.) 0 I CO = , CZ 0 0 0 7_11 r4 I a) E i .0 CC • I I ›"...1 .4 on C.O. 1 0 • .),-) co) C.) rh .--i 4 0 VI (3) ... r..., 0 a-05 171.4- ca O c/ 0 4 4 g t -r-t . o „., ..0 , —V. .. azG) 6-0)

4-1o •,...4n—I 0 ---' 90 several sections and may be excessive. The typical measured thickness is closer to 100 m (Figures 3.4, 3.5).

Lithofacies

Facies of the Alverson include basalt flows, lithic and lapilli tuffs, lapilli and bomb agglomerates, minor volcanic sandstones, volcanic plugs, and a local "crumble breccia" associated with a lava dome

(Ruisaard 1979). Flows of amygdaloidal olivine basalt constitute the

dominant facies; individual flows range from -2 m to 48 m in thick-

ness. Brecciated flow tops indicative of aa flows are typical of the basalt flows; ropy flow tops indicative of pahoehoe flows were also

reported. The fresh basalt is dark gray in color, but weathers to shades of brown, purple, and blue. Spheroidal weathering of Alverson basalt produces outrop features superficially suggestive of pillows.

Ruisaard (1979) reported pillow-like forms from one locality in Fossil

Canyon, but attributed them to subaerial pahoehoe flows rather than to subaqueous processes.

Reddish lithic (conglomeratic) and lapilli tuffs, typically unstratified and slightly to moderately welded, are also common in the

Alverson Formation. Point counts of phenocrysts suggest an andesitic to dacitic composition (Ruisaard 1979). In the lithic tuffs, the dominant clast type is volcanic, but plutonic and metamorphic clasts also occur. No evidence of marine influence has been reported from the tuffs. Lenses of agglomerate occur locally in the Alverson; one lens north of Painted Gorge attains a thickness of 72 m (Figure 3.5). A lava dome of hornblende andesite occurs at the mouth of Painted Gorge; 91 adjacent to it is a crumble breccia at the top of the Alverson (Figure 3.5). Three volcanic plugs up to 63 in in thickness occur in the CIA; they define a northwest-southeast trend (Ruisaard 1979).

Geochronology and geochemistry

Several K-Ar dates on Alverson basalt flows indicate an early to middle Miocene age (Figures 3.4-3.6). Eberly and Stanley (1978) reported a K-Ar whole-rock age of 16 + 1 Ma for a basalt flow from

Fossil (Alverson) Canyon. Ruisaard (1979) and Mace (1981) reported 7

dates on basalt flows ranging from 14.9 + 0.5 to 21.5 + 3.9 Ma. An

additional date of 24.8 + 7.4 Ma on the lava dome was reported by

Ruisaard (1979), but this age is inconsistent with the intrusive

relationship to dated flows (Figure 3.5) and is therefore discounted.

Gjerde (1982) studied geochemistry and petrography of the

Alverson basalts. He concluded that all flows from the Alverson Canyon area are alkali olivine basalts, with the exception of the uppermost

flow at Butaca Canyon (Lava Gorge), which is more tholeiitic (Figure

3.6). The alkali basalts are characterized by low alumina and high K20 and Ti02. In contrast, he found all the basalts at Painted Gorge and

Red Rock Canyon (Figures 3.5, 3.6) to be tholeiitic, generally with high alumina and low K20 and Ti02.

Ruisaard (1979) attempted to subdivide the Alverson Formation into members and to correlate these between the areas he studied. How- ever, his described sections (Figures 3.4-3.6), combined with radiometric and geochemical evidence, suggest that individual flows are 92 quite localized and that there is little lateral continuity in the internal stratigraphy of the Alverson Formation.

Lithologically equivalent units

Other occurrences of the Alverson Formation have been described from Superstition Mountain east of the Fish Creek Mountains (Dibblee

1954, 1984, Ruisaard 1979, Mace 1981, Gjerde 1982) and the Volcanic

Hills south of Sweeney Pass (Fourt 1979).

The Jacumba Basalt, named by Miller (1935) for exposures in the

Jacumba area near the Mexican border (Plate 3), is probably a southward

extension of the Alverson Formation. Hawkins (1970) and Minch and

Abbott (1973) described the Jacumba as consisting of flows of basalt

and lesser basaltic andesite, associated with agglomerates, andesitic

breccias, lahar deposits, dikes, and cinder cones. The total thickness

is typically -35 m, with a reported maximum of 75 m; individual flows

average -4 m. Geochemical data indicate that the Jacumba basalts are

high-alumina, with both alkalic and tholeiitic characteristics but with

stronger tholeiitic affinities (Hawkins 1970). K-Ar ages of 18.7 + 1.3 (hornblende) Ma (whole rock) on basalt (Hawkins 1970) and 18.5 + 0.9 Ma (Minch and and 18.6 + 0.8 (plagioclase) from an andesite breccia clast therefore Abbott 1973) indicate that the JacuMba is lower Miocene and

contemporaneous with the Alverson. On the basis of these data, the the Alverson Forma- JacuMba Basalt is virtually indistinguishable from

tion. of Barnard (1968) described a local (-5 kno) outcrop northern Baja intermediate pyroclastics in the Sierra de los Cucapas in 93

California, and named this unit the Colonia Progreso volcanics. This unit reportedly consists of up to 500 m (estimated) of hornblende andesite and dacite in massive autobreccias, interbedded with well- bedded volcanic sandstones. Inclusions of basement-derived clasts are common in some strata. The Colonia Progreso volcanics overlie metamorphic basement, grade laterally into volcanic conglomerate, and are

overlain by conglomerate with igneous and metamorphic clasts. A K-Ar

age of 15.3 + 0.8 Ma was obtained on mixed biotite and hornblende from

a dacite clast from a breccia bed (Barnard 1968); this volcanic unit is

therefore approximately coeval with the Alverson and Jacumba volcanics.

Elephant Trees member (Me) (new name)

This member comprises an assortment of thick boulder

fanglomerates and a "boulder bed" or "megabreccia", all apparently of

nonmarine origin. The Elephant Trees member is exposed in (1) Split

Mountain, where three submembers have been tentatively differentiated; Elephant (2) adjacent hills including the U. S. Gypsum property and the

Trees area, where it is left undifferentiated; and (3) the CIA between

Boundary Mountain and Red Rock Canyon (Figures 1.6, 3.1, Plate 4). The exclusive in Elephant Trees member and Alverson Formation are mutually area between geographic distribution (Figure 3.1), except for the of boulder Boundary Mountain and Red Rock Canyon, where a distal tongue

conglomerate overlies a distal tongue of the Alverson (Figure 3.6, occupy an analogous Ruisaard 1979). The Elephant Trees and Alverson crystalline base- stratigraphic position: above the Red Rock member or Imperial Formation. ment, and below the lowest marine beds of the 94 Formerly, the Elephant Trees member has been included in the Split Mountain Formation, Anza Formation, or both (Figure 1.9). This

member contains both predominantly reddish beds, generally lower in the section, and predominantly greenish-gray or olive-gray beds higher in the section. Previous workers (e. g. ver Planck 1952, Smith 1962) placed the former within a "red member" along with the Red Rock member

of this study, and the latter within a "gray member" (Figure 1.9). The

color change is sharp and obvious in Split Mountain Gorge, but in other

areas such as northwestern Split Mountain the color change is very

gradational, does not follow a major lithologic boundary, and is not

readily mappable. For example, the fanglomerates exposed in No Return

Canyon are lithologically similar to and appear to be stratigraphically

continuous with fanglomerates in Split Mountain Gorge, but the former are predominantly greenish with patches of reddish coloration, while

the latter are predominantly reddish. The color differences are probably diagenetic in origin and their stratigraphic significance is dubious.

Lower fanglomerate unit (Mel), Split Mountain

A thick, resistant, mostly monomictic boulder fanglomerate forms the northeast-facing escarpment of Split Mountain and the steep walls of most of Split Mountain gorge. This unit is equivalent to

Woodard's (1963, 1974) upper member of the Anza Formation and Kerr's

(1982, 1984) "alluvial fan facies" of the Anza Formation (Figure

1.9). Kerr (1982) reported a maximum thickness for this unit of -450 in just west of Split Mountain Gorge, and a thickness of -360 in on the 95 eastern flank of the Vallecitos. Just east of the Gorge, the thickness decreases to -150 in (Kerr 1982), to 70-75 m in Cairn Canyon (this study), and to -30 m maximum between Boundary Mountain and Red Rock

Canyon (Ruisaard 1979); it pinches out west of Red Rock Canyon. This unit rests nonconformably on basement, or with gradational contact on the Red Rock member. In southeastern Split Mountain and in Split Moun- tain Gorge it is overlain with sharp contact by the Lower Boulder Bed; in northwestern Split Mountain it is overlain with gradational contact by the upper fanglomerate unit ("Pliocene continental" facies of Kerr

1982). Kerr described the intervening contact as disconformable, but the author was unable to find evidence for an unconformity.

Kerr (1982, 1984) made a detailed sedimentologic study of this unit in the vicinity of Split Mountain Gorge, where it is particularly well exposed and accessible. He recognized three main bedding

types within the unit:

1. Debris-flow beds. These are massive, matrix-supported, boulder to cobble conglomerate beds, ranging in thickness from several cm to 12 m, with lenticular or wedge geometries. Colors range from dark red (10 R - 5 YR 4/2-3), indicative of a muddy matrix, to light red for sandy matrix. Normal, reverse, and symmetric grading were observed. In some beds, boulders protrude above the upper bedding surface.

2. Sheet-flood beds. These are generally lighter colored (10 YR

7/2), sandier, and less conglomeratic than the debris-flow beds; they range from 0.15 to 0.45 in in thickness, with an average -0.6 m.

Individual beds typically fine upward from basal granular or pebbly 96 sandstone, to low-angle cross-stratified, coarse-grained sandstone, to parallel-laminated or massive medium-grained sandstone; biotite-rich laminae are common and increase upward in individual beds.

3. Clast-supported conglomerate beds. These occur in lenticular beds up to 3 m thick; they are typically poorly stratified and dominated by cobbles and boulders.

Kerr (1982, 1984) and Kerr et al. (1979) described a systematic up-section variation in bedding types, although no logged sections were shown. The lower part of the section is reportedly dominated by sheetflood and debris-flow beds of subequal thickness, with maximum bedding thickness of 1.3 m and maximum clast size of -1 m. Farther up-section, the debris-flow beds are substantially thicker, up to 12 m, with clasts up to 2 m in diameter, whereas the sheet-flood deposits average -1 m thick. Near the top of the section, sheet-flood deposits are rare, but lenticular, clast-supported conglomerate beds are interbedded with the debris-flow beds; maximum clast size is -4 m.

Paleocurrent directions were inferred by Kerr (1982, 1984) from trough foresets, gravel A-axis bearings, channel axes, and assorted lineations, which suggest a dominant east to east-northeast transport direction (Figure 3.7). Gravel clasts are predominantly quartz diorite, with minor "spotted" tonalite, pegmatite, and biotite-quartz schist; this compos ition is similar to that of plutonic basement exposed in the

Vallecito Mountains, except that schist is not exposed in the

Vallecitos. 97

Split Mountain

foresets

gravel a-axis bearings

grooves, striations, lineations

channel axes

west of Red Rock Canyon

Figure 3.7. Paleocurrent roses for Elephant Trees member, lower fanglomerate unit (nomenclature of present study), from Kerr (1982). 98 In No Return Canyon, the lower fanglomerate unit is somewhat different in appearance, primarily in that greenish gray rather than

red is the dominant color; reddish coloration is patchy. Thick debris- flow beds, some in excess of 10 m thick, with medium-to-gray green matrix, dominate this exposure. Laminated sheetflood sandstones and

lenticular, clast-supported conglomerates also occur. Clast composition

is monomictic, similar to that in the Split Mountain exposure. In the

upper part of the canyon, the fanglamerate unit is nearly massive, with

sparse evidence of bedding, and is apparently in fault contact with

highly fractured and brecciated basement. This contact is difficult to

map because of extensive colluvial cover and the presence of a younger,

probably post-depositional fault. However, the total thickness of

fanglamerate is apparently much thinner west of the inferred fault than

to the east, and the contact is interpreted as a syndepositional normal

fault. The orientation of this inferred fault could not be determined.

Kerr (1982) interpreted the Vallecito fault, farther to the west, as

the main syndepositional fault. However, the author observed this fault

to be down-to-west, and did not observe the lower fanglomerate unit to

extend as far west as the Vallecito fault.

In Cairn Canyon, the fanglomerate unit is thinner-bedded, with

some debris-flow beds 1 to 2 m thick, but most bedding is thinner (<0.5 m), and dominated by crudely laminated conglomeratic (pebble to cobble)

sandstones. It is also thinner overall (-75 m). Clast composition is more variable than at Split Mountain Gorge, with several percent gneiss clasts. 99 Upper fanglomerate-sandstone unit (Meu), Split Mountain

In Split Mountain northwest of the gorge, a second conglomer-

atic unit overlies the first (Figure 3.1). Although it has not been

examined as closely as the lower unit just described, a few important

differences are apparent: (1) it is thinner bedded and less resistant

than the lower fanglomerate; the intervening contact forms a signifi-

cant break in slope; (2) greenish and tan beds predominate over red

beds; (3) it is more polymictic than the lower fanglomerate, and con-

tains clasts of gneiss and minor marble and amphibolite, in addition to

the ubiquitous granitoid clasts; and (4) it grades laterally toward the

southeast into marine strata. Between Oyster Shell Canyon *1 and Split

Mountain Gorge, it grades laterally into predominantly greenish, coarse

sandstone with little gravel, relatively thin bedding, abundant debris-

flow beds, sparse reddish beds, and a few thin stringers of gypsum. It

is apparently laterally equivalent to the transitional unit (MP1t) of

the Lycium member (Imperial Formation) exposed on the eastern side of

the Gorge, and possibly to the Fish Creek Gypsum as well. Placement of

the lateral boundary between this upper fanglamerate-sandstone unit of

the Elephant Trees and the lower transitional unit of the Lycium member

is arbitrary; it is placed here at Split Mountain Gorge. Stratigraphie

relations in the vicinity of Split Mountain Gorge are discussed further

in Chapter 4.

Other fanglamerates

Other fanglamerate sequences are exposed (1) in Split Mountain

on the northern flank of the Split Mountain Anticline, in the vicinity 100 of the Fish Creek Primitive Camp; (2) in an unnamed range of hills between the Elephant Trees area and the U. S. Gypsum quarry; and (3) on the U. S. Gypsum property, where it ranges in thickness from 250 m (Ver

Planck 1952) to a feather edge. In these exposures, greenish gray,

olive, or gray are the dominant colors rather than red, and ver Planck

(1952) included them in the "gray member" of the Split Mountain Forma-

tion, but stratigraphic relationships with the lower and upper

fanglomerate units of Split Mountain is unclear. Both monomictic (near

Elephant Trees) and polymictic (U. S. Gypsum property) sequences were

Observed.

In the U. S. Gypsum property and in Split Mountain east of the

Fish Creek Primitive Camp (north limb of Split Mountain Anticline),

these fanglomerates grade upward near the top into green, arkosic

sandstones with ripple lamination and debris-flow beds; the sandstones

interfinger and alternate with gypsum beds at the base of the Fish "Pliocene Creek Gypsum. Kerr (1982) designated these fanglomerates as

continental", but in the absence of direct age control or even

laterally persistent stratigraphy the age interpretation appears

unjustified.

Lower Boulder Bed (lb) ledge- The LBB is a thick, massive, olive gray (5 Y 5-6/1-2), Mountain east of forming bed that caps the NE-facing cliffs of Split single bed of the Gorge (Figure 3.1). This unit is an anomalously thick thickens matrix-supported boulders; it is typically -50 in thick, but Mountains (Kerr abruptly to -180 m against basement of the Fish Creek 101 1982). The LBB rests with sharp contact on the lower fanglamerate unit of the Elephant Trees member. In Split Mountain Gorge, -20 in of the underlying fanglomerate unit is chaotically deformed. The LBB is directly overlain over most of its extent by the Fish Creek Gypsum, although small lenses of sandstone and mudstone intervene locally. In

Split Mountain Gorge, the LBB is overlain by the upper fanglomerate- sandstone member of the Elephant Trees (west wall) or by the lower transitional unit of the Lycium member (east wall). The LBB crops out only along Split Mountain south of the Split Mountain Anticline; it pinches out to the northwest rather abruptly in the west wall of Split

Mountain Gorge, and pinches out to the southeast between Boundary

Mountain and Red Rock Canyon. It is absent north of the Split Mountain

Anticline and in the vicinity of the U. S. Gypsum quarry, where the

Fish Creek Gypsum rests directly on undifferentiated Elephant Trees fanglomerate.

Kerr (1982, 1984) and Kerr et al (1979) described details of the LBB. They characterized the unit as a megabreccia, in which a complete gradation from large boulders to sand and silt-size clasts exists; clay minerals are essentially absent. The LBB is essentially monomictic, of granitoid composition; the largest reported clast is 4.5 in in diameter. The lower contact is reportedly undulatory but concordant, and the upper contact is hummocky. At Split Mountain Gorge, the LBB is inversely graded, with the largest clasts concentrated near the top. Kerr (1982, 1984) described seams of clay and gouge-like crushed rock within the LBB; these seams are imbricated and dip toward 102 the west. Deformation of the clay seams was used to infer a southeasterly transport direction, consistent with the inferred source area in the Vallecito Mountains.

Sequence in Carrizo Impact Area.

A conglomeratic sequence occurs locally between Boundary

Mountain and Red Rock Canyon in the CIA (Figures 3.1, 3.6) and was described by Ruisaard (1979) and Kerr (1982); it has not been seen by the author. It reportedly consists of two units: (1) a lower, crudely bedded sequence of yellow to green arkosic sandstone and volcanic conglomerate, up to 30 m thick, and (2) an upper unit of plutonic-clast megabreccia, similar to the LBB, with a maximum thickness of -80 m.

Kerr (1982) described the lower unit near Red Rock Canyon as a narrow wedge of alluvial-fan deposits, with 95% basalt clasts and 5% local basement clasts, in a volcanic-lithic arkose matrix. Ruisaard

(1979) described the same unit as a debris-flow deposit. Paleocurrent directions estimated by Kerr (1982) mainly from gravel A-axis bearings are predominantly to the east (Figure 3.7), similar to results from the lower fanglamerate unit in Split Mountain.

The upper unit is similar to the LBB but was regarded by Kerr

(1982) as a separate unit because of differences in clast composition and inferred transport direction. Clasts consist of subequal amounts of granitoids, pegmatite, and schist, similar to nearby basement in the

Fish Creek Mountains. Imbricate, deformed clay seams suggest westward transport, in contrast to the southeasterly direction inferred for the

LBB (Kerr 1982). 103 Lithologically equivalent units

The Coachella Fanglomerate exposed near Whitewater Canyon nonconformably overlies basement and is overlain with angular unconformity by the marine Imperial Formation (Proctor 1968, Peterson 1975). It is also in normal-fault contact with basement; the Imperial

Formation overlaps this fault, which suggests that the fault is synde-

positional with the Coachella. A basalt flow in the Coachella has been

dated at 10 Ma (Peterson 1975). Distinctive clasts of magnetite indi-

cate that the Coachella was derived, at least in part, from the Cargo

Muchacho Mountains (Peterson 1975) or southern Chocolate Mountains (Dillon 1975).

Barnard (1968) described a unit of "older fanglomerates" in the

Sierra de los Cucapas in northernmost Baja California. This unit, up to

400 in in thickness (estimated), reportedly rests on igneous and meta-

morphic basement and consists of basement-derived clasts, except

locally where it grades laterally into volcanic conglomerate which in

turn grades laterally into the Colonia Progreso volcanics (described earlier). The "older fanglomerates" interfinger with the overlying marine Imperial Formation. Texture, clast size and composition of the fanglomerates are highly variable, and composition is closely related

in most cases to lithology of nearby basement exposures. Local anomalies where clast composition is different from nearby basement were noted by Barnard (1968), similar to those that occur in the Elephant

Trees member. 104 Fish Creek Gypsum (Hf) The Fish Creek Gypsum is a lenticular deposit of remarkably pure gypsum that crops out in Split Mountain between Split Mountain

Gorge and Boundary Mountain, north of the Split Mountain Anticline in isolated exposures, and in the U. S. Gypsum property (Figures 1.6, 3.1,

Plate 4). The gypsum is white in fresh exposures, typically with darker color banding, and weathers to a very distinctive light tan (10 YR

6-8/1). Gypsum and anhydrite occur in the subsurface below -18 m of alluvium in the anticlinal valley between the quarry and Split Moun- tain; the gypsum-anydrite transition is -50 m below the surface (Greg

Cravener, U. S. Gypsum Project Engineer, 25 April 1985, oral communication). A number of thicknesses have been reported, including (1) at

least 30 m (ver Planck 1952); (2) -66 m maximum (Woodard 1963); (3) -35 m maximum (Pappajohn 1980, Kerr et al. 1979); and (4) -60 m in the quarried area (G. Cravener, 1985, oral communication).

Along the crest of Split Mountain east of the Gorge, the gypsum rests with sharp contact on the LBB and is overlain by the lower turbidite unit of the Lycium; no interbedding with terrigenous sediment was observed. North of the Split Mountain Anticline and in the U. S.

Gypsum property, where the LBB is absent, the lower part of the gypsum is interbedded with ripple-laminated green micaceous sandstone and siltstone over a vertical distance of a few meters. However, interbedding of gypsum with demonstrably marine strata of the Imperial

Formation, as reported by Ver Planck (1952), Kerr et al. (1979), and

Pappsjohn (1980) was not observed. Interfingering small lenses of 105 gypsum in the upper Elephant Trees member can be observed at the mouth of Split Mountain Gorge across the wash from the Fish Creek Primitive

Camp, and locally in Split Mountain just west of the Gorge. Thin beds of gypsum have also been reported from the Imperial Formation in the

Coyote Mountains (Ruisaard 1979, Bell 1980); these are discussed under "Latrania member" in Chapter 4.

Bedding within the gypsum is defined by color banding, due to minor impurities of iron and manganese oxide (Ver Planck 1952). Micro- scopic grains of detrital biotite are sparsely but uniformly dispersed in the gypsum (Ver Planck 1952). In general, however, the gypsum is quite pure and free of interbedded terrigenous sediment. Gypsum is typically -95% pure in the U. S. Gypsum quarry, and is -99% pure in the area of active quarrying (G. Cravener, 1985, oral communication). CHAPTER 4

STRATIGRAPHY OF THE MARINE UNITS

The marine units include all of the Imperial Formation and part of the Split Mountain Formation of previous workers. Stratigraphic

subdivision at the formation level is too coarse to be useful for

detailed mapping and description, so the marine sequence was subdivided

into ten informal members (Figure 1.7). An attempt was made to follow

the unit boundaries and nomenclature of Woodard (1963), but a number of

changes proved to be necessary. Physical stratigraphic relationships

among the marine units are summarized in Figures 1.6, 3.1, and 4.1. Field recognition of two distinct suites of sandstone provides (1) "L"-suite a convenient basis for subdividing the marine units into

units, which are generally arkosic, coarse-grained to conglomeratic, Stone Wash, and resistant to erosion (Lycium, Latrania, Kamia Creek, units, which and Jackson Fork members), and (2) predominantly "C"-suite heavily are fine-grained, with abundant mudstones, easily eroded, and Lavender weathered in most areas (Wind Caves, Mud Hills, Deguynos, a Canyon, and Camels Head members). The Latrania member includes for mapping significant basal nonmarine facies in many localities; into purposes it is convenient to combine nonmarine and marine facies basis of one unit. The marine units can also be subdivided on the "shallow" units, sedimentologic and paleontologic criteria into: (1) 106 107

A. SOUTH LIMB, SPLIT MOUNTAIN AN tongues and intrusions of ub 7

B. NORTH LIMB MPiu MPI SE NW Mr — -\ Me 0 I — Me 0 gradational interbedded contact with interbedded gypsum gypsum & clastics

Figure 4.1. Details of stratigraphic relationships in and adjacent to Split Moutain Gorge. Not to scale. 108 where physical sedimentary structures indicate transport mainly by traction, and which typically contain shallow marine macrofauna

(Latrania, Mud Hills, Deguynos, Lavender Canyon, Camels Head, and

Jackson Forks members), and (2) "deep" units, where sedimentary structures indicate transport mainly by density currents, especially turbidity currents, and in which macrofauna are sparse to absent, although the ichnofauna may be rich (Lycium and Wind Caves members). The Stone Wash member contains facies of both kinds, but is

most conveniently mapped as a single unit. Although the "deep" units

were probably deposited in deeper than the "shallow" units, the

presence of turbidites and other sediment gravity-flow facies is not a

reliable indicator of absolute water depth, as discussed in Chapter 6.

Latrania member (MPla)

"L"-suite sandstones and conglomerates with diverse marine

macrofossil assemblages and extensive bioturbation characterize the Y 7-8/1-3), but Latrania member. The typical color is grayish yellow (5

colors range to gray or white, locally to reddish near the base tuffs), or (similar to the colors of Red Rock sandstones and Alverson sandstones to greenish where dominated by volcanic-lithic or tuffaceous and range in near the base. Sandstones are fine to coarse-grained

composition from arkose to volcanic litharenite to tuffaceous sandstone oral communica- (Ruisaard 1979, Bell 1980, Carrie Campen-Bell 1984, marble), plutonic tion). Conglomerates contain metamorphic (gneiss and basalt clasts (granitoid), and basalt clasts in varying proportions; but are the sole clast type in are absent from some conglomerate beds 109 others (Christensen 1957, Ruisaard 1979, Bell 1980). This unit contains the classic fossil localities of Fossil (Alverson) Canyon, Barrett

Canyon, and adjacent areas in the Coyote Mountains and Carrizo Impact

Area, and is probably the most intensively studied unit of the Imperial

Formation. The Latrania member does not occur in the Fish Creek-

Vallecito section per se, and was not examined in detail in the present

study. The following description of this unit is based largely on the

work of Christensen (1957), Stump (1972), Ruisaard (1979), Bell (1980),

Bell-Countryman (1984), Susan M. Kidwell (1984-1986, oral and written

communications), and Carrie Campen-Bell (1984, oral communication).

As used in this study, the name Latrania applies to the lower,

mostly marine, fossiliferous, "L"-suite member of the Imperial Forma-

tion which crops out in the Coyote Montains and on the southwestern

flank of the Fish Creek Mountains (CIA). It excludes turbiditic units and in Split Mountain (Lycium member) and the CIA (Kamia Creek member), (Stone Wash the locally fossiliferous conglomerates of Split Mountain with Hanna's (1926) ori- member). This usage is reasonably consistent sandstone unit at ginal use of "Latrania sands" for the fossiliferous areas, although no Fossil (Alverson) Canyon (Figure 4.2) and adjacent a mystery). Stump type section was given (the derivation of the name is usage to include (1972) revived the use of "Latrania" and expanded its Imperial Forma- all marine sandstones and conglomerates of the basal and "C"-suite (Wind Caves tion, including the "L"-suite (Lycium member) section (Figure 1.9). Unfor- member) turbidites of the Split Mountain

tunately, this broader definition resulted in a rather confusing

▪

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00 10'183041 00 00310 NVS "N\s‘'\\*...- 111 stratigraphic description and failure to recognize critical facies differences within the lower Imperial Formation. Christensen (1957) referred to the lowest unit of the Imperial in the Coyote Mountains as the Andrade member, and recognized three submembers: (a) a basal unfossiliferous conglomerate dominated by plutonic and metamorphic clasts; (b) a basal, unfossiliferous volcanic-clast conglomerate which locally replaces submember a; and (c) fossiliferous sandstones and conglomerates (Figure 1.8). According to Ruisaard (1979) the basal nonmarine facies range in thickness from 0 to 56 m.

Inclusion within the Latrania of unfossiliferous, probably nonmarine facies is somewhat at odds with Woodring's (1932) concept of the Imperial Formation as a marine unit. However, it makes sense for mapping purposes, as fossiliferous and bioturbated facies are inti- mately associated with unfossiliferous and non-bioturbated facies; they grade vertically and laterally and interfinger over short distances.

Furthermore, the differences between marine and nonmarine conglomerates can be subtle or cryptic.

In most areas the Latrania member rests directly on Alverson volcanics, but locally, as in Andrade Canyon (Figure 4.2), it rests on metamorphic basement. Ruisaard (1979) reported that the Latrania also rests locally on the "Split Mountain Formation" (Red Rock member of this study), but this has not been confirmed by the author. The basal contact is highly variable. Where the Latrania rests on Alverson volcanics, as in Fossil Canyon, the contact is typically an angular (up to

200) discordance (Ruisaard 1979); elsewhere the contact may be 112 conformable. However, Woodring's (1932) report of fossiliferous sandstone interbedded with basalt flows has not been supported by later investigators. Woodring (1932) also reported interstitial fossils in brecciated basalt at the top of the Alverson. In Andrade Canyon, the

Latrania member laps out against metamorphic basement in a steep nonconformity (buttress unconformity). The Latrania member probably grades laterally into the Hamia Creek and Lycium members (Figure 3.1), but the intervening contacts are covered. Claystone of the Mud Hills member

overlies the Latrania, probably with sharp contact, but the contact is

generally covered or faulted out. On the northern flank of the western

Coyote Mountains, east of Canyon Sin Nombre (Plate 4), a thin (2-3 m)

section of fossiliferous (Ostrea and Anomia), conglomeratic Latrania

sandstone with basalt pebbles and cobbles occurs between the Alverson

Formation and Camels Head member (Figure 1.6), but the nature of this

contact is somewhat Obscured by faulting.

Lithofacies and thickness of the Latrania member are highly Barrett variable (Figure 4.3). Thicknesses range from <1 m locally in central Canyon to a maximum reported thickess of -70 m in the north short Coyote Mountains (Christensen 1957). Facies variations over sections from distances are illustrated by Bell's (1980) measured

Fossil Canyon and vicinity (Figure 4.3). Few generalizations can be and upward- made about the vertical sequence, and both upward-fining relatively rare coasening trends are observed. Physical structures are

or poorly expressed, due in large part to extensive bioturbation. Bell it is not (1980) noted graded and rhythmite bedding (Figure 4.3), but 113

c, . • 1, . 0 ti " 4 n • 0.. o • - , ° 0 • 4 I 0 0 * °, 0. '0 .0 :: 00 I-0 j • • ''' o° 0. 0 0 t° 00 ** • :* ... 3 .. 000 0 • • 0 % ° .:.. p*•,., to 1: . 1 n 0 . • 0 0 0 •, .' .0 • • : • I I 114 clear from her discussion whether these are similar to graded and rhythmite bedding observed elsewhere in the Imperial Formation. Bell (1980) also described primary gypsum interbedded with siltstones in the Latrania member at Gypsum Canyon, and gypsum intraclasts in Latrania sandstone south of Tunnel Road (fig. 4.3). The gypsum beds occur be-

tween the basal, apparently nonmarine facies of the Latrania and the

overlying marine facies. Thus, their stratigraphic context is analogous

to that of the Fish Creek Gypsum, although the nomenclature of the

bracketing units is different and the two gypsum deposits are not necessarily contemporaneous.

Proximality and lapout directions can be inferred from various

criteria, including (1) onlap direction at a basal unconformity, as in

Andrade Canyon; (2) lateral facies change from sandstone to conglom-

erate, or increase in conglomerate percentage, as in lower Fossil

Canyon; and (3) thinning direction of basal conglomeratic facies (this

criterion may be ambiguous, however). These are generally apparent directions only, as outcrops rarely afford three-dimensional exposure.

Apparent proximality directions in the eastern Coyote Mountains were worked out by S. M. Kidwell (1985, oral communication), based on direct observation and inference from cross sections correlated using fossiliferous marker horizons. Proximality directions are predominantly to the west in the eastern Coyote Mountains, and more to the south in the central Coyote Mountains (Figure 4.2). This result is consistent with observations from the FCV section, where local, conglomeratic, 115

"L"-suite units (Stone Wash, Jackson Fork, Canebrake) lap out toward the west against basement of the Vallecito Mountains (Figure 2.10).

Paleontology

Macrofaunal assemblages of the basal Imperial have been discussed at length in a number of papers, including Mendenhall (1910),

Kew (1914), Vaughan (1917), Hanna (1926), Woodring (1932), Christensen

(1957), Stump (1972), Foster (1979), Bell (1980), and Bell-Countryman

(1984); Woodard (1963) reviewed the earlier studies. Several faunal

elements are particularly characteristic of the basal Imperial or

Latrania member: (1) large oysters (Ostrea heermanni Conrad) with

individual valves up to 20 cm in diameter and 4 cm in thickness; (2) a

diverse assortment of pectinids (e.g. Stump 1972) and other pelecypods;

(3) the echinoids Clypeaster (sea biscuit) and Encope (sand dollars);

(4) coral heads, both in situ as at Tunnel Road (Bell 1980) and Barrett

Canyon (Foster 1979), and as bioclasts in cobble and boulder conglom-

erates; (5) a variety of gastropods, including large (up to 15 cm

long), flattened (presumably by compaction) Conus in Barrett Canyon;

and (6) lower valves of small Ostrea and basal plates of barnacles

(Balanus) attached to cobbles and boulders. Concentrated vermiculariid

tubes form a distinctive horizon in lower Fossil Canyon. fossils are not Despite extensive bioturbation, discrete trace in abundant in the Latrania member. Ophiomorpha burrows were observed

Latrania sandstone in Fossil Canyon. Bell-Countryman (1984) reported In the large horizontal burrows attributed to spatangoid heart urchins.

Coyote Mountains, lithophagid borings are common in marble clasts 116 within Latrania conglomerates and in subcrops of marble (part of the metamorphic basement) in contact with the Latrania member (Bell 1980; S. M. Kidwell 1985, oral communication).

Lithologically equivalent units

The Stone Wash and Jackson Fork members (described later in this chapter) on the southeastern flank of the Vallecito Mountains

(Figures 1.5, 1.6, 1.12, 3.1) are locally similar in lithofacies and

biofacies to the Latrania member and could arguably be included in the

Latrania. Dibblee (1984) reported a local exposure of basal Imperial on

biotite schist at Squaw Peak north of Ocotillo Wells (Borrego Mountain

7.5' quadrangle), but did not include a detailed description.

Neogene marine strata with similar shallow-water macrofauna and

locally-derived sandstones and conglomerates have been described from

the San Gorgonio Pass region (Figure 1.1), including localities at

Whitewater Canyon (Super Creek) and Lion Canyon (Vaughan 1922, Bramkamp

1935, Woodring 1932, Allen 1957, Shremp 1981, Powell 1984, 1986). The

Imperial Formation at Super Creek consists of (1) local lenses of basal

arkosic, bioturbated sandstone, sparsely conglomeratic, with a fauna

dominated by Ostrea, pectinids, barnacles, and vermetid (?) tubes, and

(2) a more extensively exposed facies of fissile, very silty shale with

a local aragonitic fauna. Powell (1986) assigned the lower unit to the

Latrania member, and reported a thickness at Super Creek of -22 m. He assigned the upper unit to the Burrobend member (nomenclature of Stump

1972) and gave a maximum thickness of -37 m. However, the Burrobend member of Stump (1972) in the FCV section corresponds to the "C"-suite 117 marine units of this study, whereas the upper unit in Super Creek is highly micaceous and fissile, which indicates "L"-suite affinities. This upper unit is lithologically distinct from other "L"-suite marine units exposed in the Salton Trough.

In the Sierra de los Cucapas, "older fanglomerates" interfinger with claystones of the overlying marine "Imperial Formation", and where the Imperial rests directly on basement it has a basal facies of fossiliferous arkose, lithic arenite, and conglomerate (Barnard 1968). In the marine conglomerates, clasts are reportedly better rounded and sorted than in the nonmarine fanglomerates, and are commonly coated with barnacle shells and annelid tubes. Many marble clasts contain

"finger-sized holes, probably made by boring clams (Barnard 1968).

Outside the Salton Trough region, some Neogene marine units of the circum-Gulf of California also have lithofacies and biofacies similar to those of the Latrania member. These units include the Salada

Formation of eastern Baja California (Heim 1922, Durham 1950, Durham and Allison 1960, Boehm 1982, 1984); the Boleo, Gloria, and Infierno

Formations in the vicinity of Santa Rosalia (Wilson 1948, 1955); the

San Marcos, Carmen, and Marquer Formations and various unnamed marine units from islands offshore from eastern Baja (Anderson 1950, Durham

1950); and unnamed marine sandstones and conglomerates of Isla Tiburcin

(Stump 1981, Smith et al. 1985). These units range in age from middle

Miocene at Isla Tiburain (Smith et al. 1985) to late Pliocene in eastern

Baja (Durham 1950, Ingle 1974), and thus are not necessarily 118 age-equivalents of the Latrania, which is probably latest Miocene to early Pliocene (Chapter 7).

Lycium member (MP1) (reassigned name) The Lycium member consists of up to 75 in of medium to thin-

bedded, mostly coarse-grained arkosic ("L" suite) sandstones, charac-

terized by graded bedding and abundant trace fossils. It is exposed

along Split Mountain (Figures 1.5, 1.6, 1.12, 3.1, 4.1), and is essen-

tially unique to that area, although a local outlier of similar facies

(Kamia Creek member) is exposed in the Carrizo Impact Area. The Lycium

is the lowest unequivocally marine unit in the FCV section, and is

probably the lateral equivalent of the the fossiliferous Latrania

member of the Coyote Mountians and CIA. The Lycium rests with sharp

contact on the Fish Creek Gypsum, the Lower Boulder Bed, or the upper

fanglomerate-sandstone unit of the Elephant Trees member, except be-

tween Split Mountain Gorge and Oyster Shell Canyon #1, where the Lycium

interfingers with and grades laterally into the Elephant Trees member

(Figure 4.1). The Lycium is overlain with angular discordance by the

Upper Boulder Bed (UBB); where the UBB is absent, it grades up into the

Stone Wash member. Disruption and deformation of the Lycium member below the UBB will be described in the discussion of the latter unit.

Nomenclature of this unit has been somewhat problematic (Figure

1.9). Tarbet and Holman (1944) and Tarbet (1951) included it in the original definition of the Split Mountain Formation, and Woodard (1963,

1974) retained it in the restricted definition of Split Mountain, referring to it as the "marine arenite member". Elsewhere in the Salton 119

Trough, the term "Split Mountain" has come to be associated with exclusively nonmarine units (e. g. Christensen 1957, Ruisaard 1979,

Dibblee 1984). Thus, in the Coyote Mountains and CIA, "Split Mountain

Formation" refers to strata underlying the Alverson volcanics

(Christensen 1957, Ruisaard 1979), whereas marine sandstones of the

type Split Mountain (referred to here as Lycium member) are strati-

graphically higher than the LBB and therefore probably higher than the

Alverson as well. Woodard (1963) previously applied the informal name

"Lycium member" to marine sandstones above the UBB (Wind Caves member

of this study). However, that unit is not well exposed in Lycium Wash,

whereas the present unit under discussion is well exposed in Lycium

Canyon; therefore, the informal name has been reassigned here.

At least three units can be recognized within the Lycium member

on the basis of differences in bedding type and geometry. They are

spatially separated but do not constitute separately mappable units at in a scale of 1:24,000; stratigraphic relationships are summarized

Figure 4.1.

Transitional unit (MP1t) east wall of This unit is a very local facies exposed in the light yellowish Split Mountain Gorge, where it consists of -21 m of olive (5 Y 5/2) thin gray (5 y 7/1), olive gray (5 Y 3/2-3), and dark below the lower to medium-bedded sandstones exposed above the LBB and is also turbidite unit of the Lycium member (Figure 4.4). This unit unit to the east, laterally transitional between the lower turbidite unit of which is clearly marine, and the upper fanglomerate-sandstone 120

ASCE -40 ABE, <-9SM-3-8? BE ACE ABE .._ BECS A B(C)E ABE

—C. ABE -35

ABE

_AC- ABE AE " BECS -30

MPW MPH BE Br/ CE -25

AE, BE, 'BCE MPH

MPft -20

.-CE ABE

/..A7ffe ABE,BE, _ -15 55 B CE CE SM -2a

A (B )E

}ABE

.. -10 50 ...... 'ABE

ABE BE 45 -5 ABE BE A E .. I ABE, BE, CE tICE ABE BE CE ABE 46

SM-2

Figure 4.4. Graphic logs of measured section SM-2, Split Mountain Gorge. Location is shown in Figure 1.12 and Plate 4. Base of section in this and all other graphic logs is at lower left; top is at upper right. Symbols are explained in Figure 1.13. This log is also shown in Plate 5. 121 the Elephant Trees member to the west, which is probably nonmarine (Figure 4.1). The most distinctive characteristic of this unit is abundant dark-olive debris-flow beds, which range from a few cm to nearly 1 m in thickness. Other sedimentary structures include ripple lamination and rare graded bedding. The top of the transitional unit is marked by -50 cm of parallel-laminated sandstone (Figure 4.4). No marine or paleocurrent indicators were observed in this unit. The equivalent strata on the west side of the gorge are similar in appearance but contain some reddish beds; they are arbitrarily included in the upper fanglomerate-sandstone unit of the Elephant Trees member (Chapter 3).

Lower (thin-bedded) turbidite unit (MP11)

This poorly exposed unit contains the lowest unequivocally marine beds in the Split Mountain section. It occurs along Split Moun- tain east of the Gorge; the best exposures are in upper Crazycline

Canyon (where a thickness of -32 m was measured), upper Wind Caves

Canyon, and Split Mountain Gorge (where it is -5 m thick - Figure

4.4). It overlies the Fish Creek Gypsum or (in the gorge) the transitional unit. It grades up into the upper turbidite unit, or where the upper turbidite unit has been truncated by the UBB, it is overlain discordantly by the UBB. In most areas the lower turbidite unit and its contacts are poorly exposed due to colluvial cover; its presence is inferred from float.

Where well exposed, as in upper Crazycline Canyon, the lower turbidite unit is characterized by thin (1-10 cm), parallel, graded 122 sandstone beds of uniform thickness. Beds are predominantly fine to medium grained, the thicker beds generally being coarser. These sandstones are yellowish gray (5 Y 8/1), with darker (5-10 Y 6-4/1-2) silty interbeds. This unit is finer-grained and more thinly and uniformly bedded than the overlying upper turbidite unit. Partial Bouma (1962) sequences, including AE, ABE, BCE (uncommon), and BE, are characteristic of most beds. Mixing by bioturbation of the upper parts of sand-

stone beds with interbedded siltstones helps to accentuate the grading, but tends to obscure primary physical structures of the Bouma

sequences. Horizontal burrows are the dominant trace fossil, including

well-preserved Scolicia. A few bidirectional sole marks, and one uni-

directional flute cast, were found in upper Wind Caves Wash; these indicate a southerly transport direction (Figure 4.5).

Upper turbidite unit (MPlu)

This unit is volumetrically dominant in the Lycium member. It

is conspicuously exposed in Split Mountain Gorge, where it is characterized by medium-bedded, lenticular, graded, coarse-grained sandstones weathered to brown (10 YR 7/2 to 5 YR 5/4). While this is the best known exposure of this unit, fresher exposures of light yellowish gray

(5 Y 7-8/1) sandstone occur in Oyster Shell and Lycium Canyons. Sec- tions were logged from all three areas (Figure 4.4, 4.6-4.9, Plate

5). This unit is also exposed east of Split Mountain Gorge, where it is chaotically deformed below the UBB. Chaotic, outcrop-scale folds in the

Upper turbidite unit are best exposed in Crazycline Canyon. Farther to 123

N N 10% 40% n=4

n =39 Stone Wash member Lycium member - upper turbidite unit

N 2C)9n)

n=10

Lycium member - lower turbidite unit

III unidirectional } sole marks bidirectional

Figure 4.5. Paleocurrent roses for "L"-suite marine units.

124

40 PC

• -35

AAE ,iRG ,••

_a- ABC

CE CE B CE ABE

BCE CE AE

. •

BLE,CE

-25 MPW

beds lent fouler- abundant CBI 8. fill

MPlu SM-3

Figure 4.6. Graphic logs of measured section SMH3, Split Mountain Gorge. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 5.

125

cg VC 111 vl SS SS SS SS SS 80 105 Pm II°

offset 0 41, ocr, -67; 1111.1 ..o°040 •0 110 7/17 jar%

,00000..• t

offset 41111. 9/ 4 II° °CIO 0° 9• -30 00••. distal n -70 105 ;*•<-70p °:0% tongue of Ps WIMP. (3* 0g7)-:• '' FC

40 :IC,: pr Zee., ED 231- g 0 cp ?,

CD,• -) -25 DD 100

c:D2,.. •

0°.1•°;%;'47°) •

4,00 -,,„ CZ) .;.:.:. 85 0c52,°',0 a 0- 906,0k:64 . n 0 ix) :. f../P,•?-:; *•°:: • z, <1 o oz

80 Wmdas of ab

offset r.X",g4;fr »0'00

Figure 4.7. Graphic logs of measured section LW-1, Lycium Canyon. - Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 126

OS-2

Figure 4.8. Graphic logs of measured section OS-2, Oyster Shell Canyon *3. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 127

Figure 4.9. Graphic logs of measured section OS-1, Oyster Shell Canyon *3. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 5. 128 the east, this unit is deeply truncated by the UBB and only locally exposed.

In Split Mountain Gorge, the main bedding types in this unit are medium graded beds, typically 20 cm to 1 in thick, and thin graded beds, typically a few cm thick (Figures 4.4, 4.6). In the medium-bedded sequences, Bouma division A is dominant, B is common, and C is uncom-

mon. D and E divisions consists of sandy siltstone with little clay;

these are mostly intensely burrowed and bioturbated and cannot be

differentiated. Associated structures include flame structures, tool

marks, groove casts, flute casts, and load casts. Rip-up clasts, mostly

of fine-grained D-E division bed tops, are common and are concentrated

in some beds as trains or clusters. Many A divisions, otherwise mas-

sive, show discontinuous, subparallel streaky laminations. These may be

similar to Lowe's (1982) "traction carpet" layers, commonly developed

in high-density turbidites. However, dewatering structures such as dish

structures, also characteristic of high-density turbidites, were not

observed. Reverse grading occurs at the base of some beds.

Thinner beds are an important secondary lithofacies in the

upper turbidite unit. These typically lack the Bouma A division, but B

and C are common, and some beds resemble wavy or lenticular bedding.

Bioturbation is particularly intense in these thinner beds. The thinner beds are typically clustered and in some outcrops appear to thicken

laterally.

In Oyster Shell and Lycium Canyons (Figures 4.7, 4.8), Bouma sequences are less well developed than in Split Mountain Gorge. Massive 129 graded beds, commonly with discontinuous, streaky laminations, are more typical. In many beds, grading may be a result primarily of bioturbation of the upper parts of sandstone beds ("bioturbation grading").

Where the upper turbidite unit grades into coarser-grained, conglomeratic facies, some ungraded beds appear. This is illustrated in Section

LW-1 from Lycium Canyon (Figure 4.7), where the Lycium member grades up into the Stone Wash member. Some of the ungraded beds are trough cross- stratified. Other occurrences are between Split Mountain Gorge and

Oyster Shell Canyon #1, where the Lycium member grades laterally into the Elephant Trees member (upper fanglomerate-sandstone unit), and in the northern part of Lycium Canyon. This last area is not well exposed, however. Small-scale scour-and-fill channels also occur in the upper part of the section. These are typically -1 in deep and <10 in wide.

A number of paleocurrent readings were made on groove casts and flute casts. Flute casts are not sufficiently exposed to give an unambiguous sense of transport directions. However, trains of rip-up clasts

traceable to the underlying bed top consistently indicate a sense of transport to the east. Paleocurrent data indicate a dominant transport direction toward the east to northeast (Figure 4.5).

Boulders and cobbles are incorporated in sandstone beds, particularly in the upper part of the section. These are commonly clustered, as in boulder nests. Boulders are typically subrounded and well sorted; the combination with sandstone matrix results in a strikingly bimodal texture. 130 The most unusual and enigmatic aspect of the upper turbidite unit is outcrop-scale sigmoid stratification. The style of stratification is most apparent in Lycium Canyon, but can also be observed in

Oyster Shell Canyons and Split Mountain Gorge. Similar stratification

geometries were also observed in the distal facies of the Stone Wash member, in "L"-suite sandstones of the Wind Caves member, and in the

Kamia Creek member. Packages of sigmoid clinoforms are -5-15 m thick,

and clinoforms accrete in the upcurrent direction, i. e. , toward the

west. If these result from migration of wavelike bedforms, the wave-

length is too long to observe in outcrop. Divergence in dip direction

between backsets and set boundaries, and angular discordances at set boundaries can be as much as 15°. This bedding type is further dis-

cussed in Chapter 6.

Paleontology and biostratigraphy

The Lycium member is the lowest marine unit in the Split Moun-

tain sequence, and possibly the oldest marine unit exposed in the

western Salton Trough. Therefore, its age is important for determining

the marine transgressive history of the Salton Trough, and

biostratigraphy provides the most direct constraint on its age. Unlike

the Latrania member, the Lycium contains essentially no macrofossils,

except for rare, reworked valves of Ostrea heermanni.

Tarbet and Holman (1944) reported that "a meager [foramin-

iferal] fauna from the Split Mountain formation suggests Miocene age."

Durham and Allison (1961) listed the foraminifers Hanzawaia, Reophax,

Textularia, Textulariella, and Globigerinoides from the same unit, and 131 stated that "it [Lycium member of this study] is clearly separated from the overlying Imperial Formation by an erosional unconformity [presum-

ably the base of the UBB]"; they also proposed a Miocene age. However,

W. H. Holman (1969), in a written communication to L. A. Tarbet, cited

by Pappajohn (1980), stated:

It seems reasonable to place these Split Mountain bedsin the Pliocene, along with the Imperial. . . . Any conceivable age difference as shown by forams in Split Mountain beds would necessarily be regarded as being much less than epoch magnitude.

The "Split Mountain beds" apparently correspond to the Lycium member as

used here. Holman (1973) later stated, in a written communication to

G. Gastil, cited by Pappajohn (1980):

Samples [of microfossils] from the Imperial Formation about ten feet above gypsum deposits at Split Mountain [half a mile east of the mouth of the gorge], yielded all the species of the Gorge Fauna and several species which also appear higher in the Imperial Formation. From this it was concluded that the foram beds concerned are part of the initial transgressive sequence of the Imperial Formation. The Miocene Age stated for the Imperial Formation [Tarbet and Holman, 1944] was used following Woodring, because the age was expressed with respect to Epochs which were originally set up by using macrofossils; evidence from forams was considered inadequate. . . . [Since then], Imperial Age has been established as Pliocene. lower The microfossil locality referred to is apparently in the in the (thin-bedded) turbidite unit of the Lycium member, probably

Crazycline Canyon exposure. Pappajohn (1980) added that: the Additional examinations of Split Mountain fauna from marine sandstone and shale above the gypsum were con- ducted by Mobil Oil Company later in 1973 (personal to G. Gastil). Their findings, based on communication for forams, indicate an Upper Miocene or younger age these beds. Age determinations based on calcareous 132 nannofossils yielded Late Miocene to Early Pliocene (zones NN10-NN16, Bramlette and Wilcoxon, 1967). According to Stump (1972), "All the samples in the lowermost Latrania

member [at Split Mountain: Lycium member of this study] contain only nondiagnostic arenaceous benthonic foraminifers." Ingle (1974) reported only "scattered shallow water microfossils" from the marine "Split Mountain Formation."

In conclusion, biostratigraphic resolution appears to be insuf-

ficient to distinguish between latest Miocene and earliest Pliocene ages

for the lowest marine beds. Further evidence for the earliest trans-

gression of the Salton Trough is discussed in Chapters 7 and 8.

Trace fossils are abundant in the Lycium member, and are best

exposed in Oyster Shell and-Lycium Canyons. Horizontal burrows predomi-

nate and are concentrated in the upper, fine-grained, darker-

colored parts of graded beds. Scolicia is the most distinctive and

identifiable of the horizontal burrows. Ophiomorpha burrows also occur

in the upper turbidite unit of the Lycium; although not as common as

Scolicia they are quite conspicuous. Thalassinoides, Chondrites, and an

unnamed "feather-duster" trace fossil were also observed in the upper

turbidite unit of the Lycium member.

Kamia Creek member (Mph) (new name)

Lithofacies of the Kamia Creek member are similar to those of

the upper turbidite unit of the Lycium member; the Kamia Creek consists of arkosic, coarse-grained, graded sandstone beds, with granule to pebble gravel and isolated boulders (Figure 4.10). This member occurs in 133

cg vc c m I VI StrSJCS os SS SS su us 80 4 t ^ 155 L Plc large Oh cartr=E=r=. offset co. •

pOps 35 -75 -150

/////

;•- pe.Os

tuffaceous 30 70 -145 large offset

-140 25 65

Plc Pm

-100 20 IMP 60

95 15 ...... 55

ASE

0 Pm 17-9° 10 50 MPk .

pe.Os 85' 5 F

... pe.Os 80 o 40 o KC-1

Figure 4.10. Graphic logs of measured section KC-1, Kamia Creek, CIA. Location is shown in Plate 4. This log is also shown in Plate 5. 134 a single exposure of west-dipping beds in the Carrizo Impact Area at the head of Kamia Creek (Figure 3.1, Plate 4). Only the upper contact with the Mud Hills member is exposed; the Kamia Creek member probably overlies the Alverson volcanics and grades laterally into the Latrania member, but those contacts are covered with alluvium.

Bedding in the Kamia Creek member is similar to that of the

Lycium; lenticular graded beds of predominantly coarse-grained sandstone with local isolated boulders make up most of the section. In many beds,

partial Bouma sequences (mainly ABE) can be seen; in other beds, biotur-

bation obscures details of the upper parts of graded beds. The graded A

division of many beds contains granule to pebble gravel. The Kamia Creek

member grades up into a sequence of fossiliferous sandstone and con-

glomeratic sandstone without graded bedding and with a shallow-water

macrofauna, dominated by pectinids and Ostrea.

Two bidirectional sole marks at the bases of graded beds indi-

cate predominantly easterly or westerly paleocurrents (Figure 4.5). -

Pebbles and cobbles are of granitoid composition; no volcanic clasts

were observed, even though the Kamia Creek apparently overlies Alverson

volcanics and the Alverson crops out nearby.

Stone Wash member (MPs) (new name) local marine Conglomerates and conglomeratic sandstones with and Stone Washes indicators crop out in the drainage basins of Lycium Wash (Figures 1.12, 3.1), and are referred to here as the Stone

member. In Split Mountain west of the gorge, the Stone Wash unit of the conglomerates are separated from the upper conglomerate 135 Elephant Trees member by intervening marine sandstones of the Lycium member; the Lycium grades up into the Stone Wash member (Figure 4.7). The Stone Wash member may merge with the Elephant Trees conglomerates in the vicinity of No Return Canyon, and with Jackson Fork member between Coral Wash and Jackson Fork, but stratigraphic relationships there are obscured by faulting and colluvial cover. West of Coral Wash the Stone Wash member laps out against granitoid basement of the Vallecito Mountains. Toward the southeast, in Oyster Shell Wash, the Stone Wash member interfingers with and grades into the lower ("L"- suite) part of the Wind Caves member (Figure 4.9). The Stone Wash member is overlain by the Mud Hills member with sharp or interfingering contact; isolated tongues of conglomerate and conglomeratic sandstone within massive claystone of the Mud Hills member are exposed along lower

Lycium Wash. The Upper Boulder Bed (UBB), which forms the dip slope of much of Split Mountain, pinches out as a tongue within the distal Stone

Wash member, and is treated here as a facies of the Stone Wash member.

Because of these complex stratigraphic relationships (Figure

4.1), this member has not been adequately described or differentiated in previous studies (Figure 1.10). Tarbet and Holman (1944) and Dibblee

(1954) apparently included it in the Split Mountain formation, and

Woodard (1963, 1974) mapped it (along with the UBB) as the "upper fanglomerate member" of the Split Mountain Formation. Woodard (1963) also mentioned "coarse arenite and pebbly conglomerate beds in the upper reaches of Lycium Wash and west of the prominent mesa [Hill

1].34?]" under the description of his Lycium member; this description 136 fits the Stone Wash member of this study. Kerr (1982) mapped this unit as "Pliocene continental beds". However, unlike the nonmarine Elephant Trees fanglomerates, the Stone Wash member exhibits local marine macro- and faunas locally intensive bioturbation of the sandstone matrix. The

Stone Wash member can be conveniently subdivided into a proximal facies, which is similar to the fossiliferous Latrania member, a distal facies, which has features indicative of subaqueous sediment-gravity flows and grades into the Lycium and Wind Caves members, and the Upper Boulder Bed, an extensive megabreccia bed.

Proximal facies

North and west of Hill 1134, the Stone Wash member contains local concentrations of macrofossils, reminiscent of the Latrania fauna but generally sparser. Bioturbation of the sandstone matrix is pervasive in some localities, similar to the Latrania member. The most characteristic fossils are (1) valves of Ostrea heermanni, Ostrea vespertina, and Anomia subcostata; (2) inflated "sea-biscuit" echinoids (Clypeaster bowersi according to Woodard 1963) up to 10 cm in diameter; (3) rare coral heads with lithophagid borings, occurring as cobble-size bioclasts within conglomerate beds; (4) lower valves of Ostrea vespertina and basal plates of Balanus attached to boulders and cobbles; and (5) pectinids. S. M. Kidwell (1986, written communication) has provided a preliminary faunal list for the Stone Wash member (Table 4.1). 137 Table 4.1. Tentative faunal list for Stone Wash member (MPs), Coral Wash exposures, based on field identifications by S. M. Kidwell (1986, written communication).

1. Associated with cobble conglomerate

coral heads attached to cobbles Ostrea heermanni lithophagid borings in coral and O. heermanni Cliona borings in O. heermanni

2. Associated with granular to coarse sandstone

Turritella imperialis (whorl diameter to 2 cm) Evola keepi (pectinid) "Conus" "Spisula" mold cardiid bivalve regular echinoid spine Clypeaster Pecten sp. Dosinia molds venerid bivalve mold Encope fragments Anadara mold Ostrea heermanni 138 Distal facies

Bedding in the distal conglomerate facies is rather poorly developed, especially in comparison with the Elephant Trees member, and is defined mainly by variations in clast size, sorting, and concentra-

tion. Matrix is typically coarse sandstone, and fabrics are mostly matrix-supported, although clast-supported fabrics also occur. In Lycium

Canyon (Figure 4.7), the conglomerate is polymictic, with abundant

gneiss boulders and cobbles, whereas along Coral Wash the conglomerate

is monomictic and virtually identical in composition to the underlying

granitoid basement. In Lycium Canyon, discontinuous partings of dark to

mottled, silty sandstones occur within otherwise poorly organized con-

glomerates. These partings are similar to bioturbated bedding tops in

the Lycium member, except that discrete burrows are generally difficult

to see. In upper Oyster Shell Canyon #3 (Figure 4.8), thick to massive

conglomerate beds alternate with graded conglomeratic sandstone beds similar to those of the Lycium and basal Wind Caves member, into which

the distal Stone Wash member grades. In all exposures the distal Stone

Wash member lacks megafossils or evidence for pervasive bioturbation of the matrix.

Upper Boulder Bed (ub)

The UBB is a thick, massive, chaotic bed of essentially unsorted clasts ranging in size up to 10 m in a dark grayish olive (5-10 Y 4/1-2) matrix. The UBB is resistant to erosion and forms the prominent southwest-facing dip slope of much of Split Mountain from Oyster Shell

Wash to Boundary Mountain. Kerr (1982) showed its easternmost outcrop to 139 be east of Boundary Mountain. It typically occurs between turbidite sandstones of the Lycium and Wind Caves members, but in the western

limit of its exposure it is essentially a bed within the Stone Wash member (Figure 4.7), and the UBB can be considered as part of the Stone

Wash member. The typical thickness of the UBB is -20-25 m and Pappajohn

(1980) reported a maximum thickness of 50 m, but thickness is difficult

to measure because of an irregular basal surface, which includes large

protrusions into the underlying Lycium turbidites. The UBB is similar to

the LBB in terms of texture, outcrop appearance, scale, and geographic

distribution. However, the stratigraphic context of the UBB indicates

submarine emplacement, whereas the LBB is probably subaerial. The UBB is

also very polymictic, with abundant gneiss clasts. Local monomictic

concentrations of angular clasts are common, including light-colored

streaks of pegmatite fragments; these suggest brecciation during trans-

port. In places, as at the head of Split Mountain Gorge, the UBB grades

upward into sandstone.

Disruption of the Lycium turbidites below the UBB is

spectacular, especially in Split Mountain Gorge and Crazycline Gorge at Canyon. The best-known example is exposed in Split Mountain Lycium mem- wash level, where a dipping tongue of the UBB intrudes the tongue are ber, and the turbidite beds on the downdip side of the tongue of UBB deformed into disharmonic folds. Farther up the gorge a north (perhaps an extension of the same one) is exposed high on the underlying bedding can be seen, wall. In this exposure the continuity of

demonstrating that the structure is not fault-related. Numerous 140 disharmonic folds on the scale of several meters are exposed in

Crazycline Canyon. Axial orientations were measured on several folds

(Figure 4.11); these show a preference for northeast-southwest

orientations, but no preferential sense of overturning was seen.

Inclusions of large (up to several meters) rip-up flaps of

turbidite sandstone are common in the UBB (Figure 4.8). In Oyster Shell

Canyon #2, inclusions of conglomerate beds were also observed in the

UBB. A particularly enigmatic lithology is exposed in Oyster Shell

Canyon #2 and Lycium Canyon; it consists of -1 m inclusions of the UBB

lithology enclosed in a lighter-colored conglomerate bed. This bed

appears to overlie the UBB.

Wind Caves member (Pw) (new name)

This member comprises approximately 200 m of thin to very

thick-bedded, friable, yellow-gray (5 Y 7/2) sandstone, mostly fine to

medium-grained, interbedded with subordinate gray claystones. Sand-

stones are characterized by graded bedding, sole marks, and partial member, Bouma sequences, as in the Lycium member. Unlike the Lycium "L"- however, sandstones are predominantly of the "C" suite, although all the suite sandstone beds occur at and near the base of the unit in "C" sand- measured sections (Figures 4.12-4.17), and alternate with

stones throughout section NF-1 (Figure 4.12). southwest This member is named after the "Wind Caves" on the Cames member facing slope of Split Mountain east of the gorge. The Wind

of this study is essentially the same as Woodards's (1963) Lycium

member; however, that name was reassigned to the lower turbidites 141

Figure 4.11. Stereonet plot (lower hemisphere projection) of axes of outcrop-scale folds in Lycium member below Upper Boulder Bed. Arrows indicate sense of overturning. Attitudes are not corrected for regional structural dip, but average bedding for Split Mountain Gorge area is shown. SMG, Split Mountain Gorge; 0S1, Oyster Shell Canyon *1; all others are in Crazycline Canyon.

142

Sc rn cg ss $S $ s SS Pw 115

-35 -75

ACE

AfgE 110 some ACE thin ACE • •

beds -30 -70 ACE

CE ACE BE ACE -105 C? AEACE • h

BCE 25 -65 ACEcE_...r CEABE •C ABCE-4- PL ABLE BCE( ABE ,f - — 100 ABCE-t." A- s AB BEACE

ABCE 20 -60 BCE-. ABE C

BE/BCE BE C,_55 BE 15

F 10 -50

-85

5 -45 p. ABE

trtfij,- BCE p. 80 CE .. Bof •• .. large pw offset ▪ L_o 40

f / top of ub short distance below; not exposed NF-1 (part)

Figure 4.12. Graphic log of lower part of measured section NF-1, North Fork. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 143

Cg VC of SS so L 165

ACE IICAEC{e -205 -245

v-L

160 ACE

-200 Pw A -240

155

k-oc1C 195 -235 —71 L

ABCE/

150 O ftet L -190 -230 ACE BCE CE

145 -185 -225 CE, BCE, ACE

v-Th -140 L -220 180 C&—L

v-L k-Th structurally/ "= disrupted ACE FC -195 -r- •••• -215 CE ABE,

StmOmaIly disrupted P//'

C-2W 11,1

ACE

Figure 4.13. Graphic log of upper part of measured section NF-1, North Fork. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 144

115

110

105

100

SM-1

Figure 4.14. Graphic log of measured section SM-1, head of Split Mountain Gorge. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 145

cg ivc In vi splçsj so ss ss ss ss

PwH-

65 105

60 100

95 55 , L

50 90 large offset n/7

45 85

40 A 80

Figure 4.15. Graphic log of measured section FC-5, Fish Creek Wash near head of Split Mountain Gorge. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 146

Cg vc1 c ml I lvf strs.lc.sj ss ss ss ss ss 80 Pw "i7

•• • ACE-.

-65

-60

-55

-50

up to 5 m of cs., cgtc. turbldites -45 fill paleorenet on upper contact of ub

Of f8 C 40

EK-1 (part)

Figure 4.16. Graphic log of lower part of measured section EK-1, Cairn Canyon. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 147

Cg vci cim f vl us us SS SS SS 120 200 C 160 Pw

115 —se .L -166 C-196

ABCE 110 -150 -190

105 -145 -185

1.0

Offset -100 -140 -180 difficult FC correlation

-95 -135 -175

90 -130 -170 er"

offset L— difficult ACE conolation f 4-L -125 BE L 166 -85 ABCE CE ACE ACE t ACE C

ABCE 160 Pw 80 120 c

Figure 4.17. Graphic log of upper part of measured section EK-1, Cairn Canyon. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 6. 148

(underlying the Stone Wash member and UBB) which are well exposed in

Lycium Canyon. The complete Wind Caves section is well exposed only

along North Fork (Figures 4.12, 4.13) and Cairn Wash (Figures 4.16,

4.17); but good partial exposures occur along Fish Creek Wash near the

mouth of Split Mountain Gorge (Figures 4.14, 4.15). All four sections

are summarized in Plate 6. At its northwestern limit, the Wind Caves member grades later-

ally into (or interfingers with) the Stone Wash member, and probably

interfingers with the Mud Hills member (Figure 3.1). Details of these

lateral gradations are obscured by alluvial-fan terraces and colluvium;

the gradation into the Stone Wash member is best exposed along the

inappropriately named Oyster Shell Wash. At its southeastern limit, the

Wind Caves member appears to pinch out or shale out into the Mud Hills and member, but this relationship is obscured by alluvial-fan terraces

faulting along the southwestern flank of Split Mountain southeast of

Cairn Wash. The base of the Wind Caves member is defined as the top of 2 the UBB; the top is defined as the uppermost sandstone bed at least

cm thick. This uppermost contact is difficult to locate precisely

except along fresh cuts in washes.

Sedimentary facies member display a "C"-suite sandstone beds in the Wind Caves are gradation between two end members. At one extreme, sandstone beds

very thick, up to 8 m, with undulatory contacts (Figures 4.14, 4.15); combination of these beds are presumably lenticular, although the limited lateral friability, weathering, structural deformation, and 149 exposure obscures the geometry of sandstone beds. These very thick beds typically contain gravel, plus uncommon cobbles and boulders, concentrated near the base of individual beds. Most boulders are granitoid- plutonic or gneiss clasts, similar to nearby basement rocks. One boulder, exposed in lower Wind Cames Wash, consists of a large reworked cleat of the underlying Upper Boulder Bed. Very thick beds of the Wind Caves member are dominated by the Bouma A division, but B and C divisions are common as well, and convolute lamination occurs near the top of some beds. Spectacular load casts and flute casts are developed on the bases of some beds. This facies is best developed along Fish Creek

Wash just above the head of Split Mountain Wash, and is illustrated in sections SM-1 (Figure 4.14) and FC-5 (Figure 4.15). These sections represent the lower part of the Wind Caves section and the approximate center of the Wind Caves outcrop belt. In laterally adjacent exposures, at sections FC-1 (Figures 4.12, 4.13) and EK-1 (Figures 4.16, 4.17) the corresponding part of the section is characterized by thinner bedding

(typically 0.5 to 2 m).

The other extreme of "C"-suite bedding type is well illustrated in the upper part of sections FC-1 and EK-1. Relatively thin-bedded, flysch-like sandstones are laterally persistent with fairly uniform thickness at outcrop scale (a few tens of meters). Bouma sequences are well developed; AE, ABE, ACE, BCE, and ABCE Bouma sequences can all be recognized. Near the top of the section, exposed in lower Lycium Wash, thin, graded beds are locally contorted, suggestive of penecontempor- aneous deformation. 150

"L"-suite sandstone beds are extensively exposed in the basal part of the Wind Caves member, particularly in Oyster Shell Wash and

Canyons, and at the Wind Caves. These sandstones are mostly medium- bedded, graded, and lenticular, and very similar in overall appearance to the upper turbidite unit of the Lycium member. Lenticular bedding geometries suggestive of retrogradational sigmoid clinoforms, similar

to those in the Lycium member, were observed in "L"-suite sandstones of

the Wind Caves member along Oyster Shell Wash. "L"-suite sandstones also occur interbedded with "C"-suite sandstones (e. g. Figures 4.12,

4.13); these beds are generally thinner and more laterally persistent

than "L"-suite sandstones at the base of the Wind Caves member.

Paleocurrents

Paleocurrent directions were measured from a variety of sole

marks, both unidirectional (flute casts) and bidirectional (groove

casts and incompletely exposed flute casts). These paleocurrent indi- 4.18), as cators show a dominant southerly transport direction (Figure

observed by Smith (1962). Pappajohn (1980) apparently overlooked this Lycium mem- dominant mode by compositing data from the Wind Caves and

bers of this study, which have quite different modes. Caves member When subdivided geographically, data from the Wind (Figure are generally less dispersed and show two additional details

4.18). First, modal directions rotate from south-southwest to south to exposures (Oyster Shell south-southeast, from the northwesternmost other words, Wash) to the southeasternmost exposures (Cairn Wash). In from Oyster Shell Wash the dispersal pattern is radial. Second, data 151

20%

n=10

OS-m

WC-m Section NF -1 EK

Sole Marks

I I bidirectional

unidirectiopal

ripple crests

AH Data

Figure 4.18. Paleocurrent roses for Wind Caves member, subdivided by area (top) and composited (bottom). 152 show a secondary easterly mode, associated with "L" sandstone beds, reflecting the lateral gradation into the Stone Wash member which has predominantly easterly paleocurrent directions (Figure 4.5).

Paleontology

Macrofossils are rare in this unit. One valve of the thick- walled oyster Ostrea heermani was found in the North Fork section

(NF-1-16.5), probably reworked from the laterally equivalent Stone Wash member, in which such oyster shells are locally common. Woodard (1963) reported isolated valves of the thick-shelled oyster Ostrea heermanni,

and later (1974) an echinoid, an unidentified marine bivalve, and bryozoan fossils from the lower 60 m of his Imperial Formation (possi-

bly the Wind Caves member of this study). Bioturbation is less obvious

than in the Lycium member, and discrete trace fossils are rare. Gyro-

lithes and Thalassinoides were found in North Fork (NF-1-93-108, Figure

4.12); these trace fossils are more characteristic of the Deguynos

member.

In contrast to the sparse macrofauna and ichnofauna, the micro

fauna of the Wind Caves member is abundant and diverse. Stump (1972)

listed a planktonic foraminiferal fauna which he assigned to Zone N19 Globoquadrina (lower Pliocene) of Blow (1969), based on the presence of (1974) con- altispira altispira and Sphaeroidinella dehiscens. Ingle decrease in curred with Stump's age assignment, and reported an abrupt Forma- planktonic foraminifers -200 in above the base of his Imperial

tion, (i. e. near the top of the Wind Caves member). 153 The benthic microfauna, according to Ingle (1974), is predominantly subtropical and similar to that of the modern Gulf, based on the predominance of dextral-coiling "Globigerina" pachyderma. Sinistral "G." pachyderma also occurs locally, suggesting a temperate influence. Ingle assigned the benthic fauna to the upper bathyal (>100 in water

depth) to "shelf" paleobathymetric zones. Quinn and Cronin (1984) the interpreted foraminiferal assemblage from the lower part of their

"middle Imperial Formation" (apparently the Wind Caves member of this

study) as "outer shelf". Principal component analysis of foraminiferal

assemblages showed this unit (their samples 15 and 16) to be signifi-

cantly different from other units of the Imperial Formation.

Lithologically equivalent units

No lithologic equivalents of the Wind Caves member are known to

crop out elsewhere in the Salton Trough region. However, an unnamed sequence of marine sandstones and mudstones with similar character-

istics has been reported from the Fortuna and San Luis Basins of the

Yuma Desert, referred to as the "older sedimentary rocks" of ques- tionable Miocene age (Olmsted et al. 1973) and as "marine upper Miocene strata" (Eberly and Stanley 1978). According to these reports, dipmeter data suggest angular discordances between this unnamed marine unit and the overlying marine Bouse Formation, and between it and underlying coarse terrigenous sediment. For this reason, the unnamed marine unit has been tentatively assigned by some to the Miocene, despite the

Presence of Sphaeroidinella dehiscens (Olmsted et al. 1973, Winterer

1975), commonly regarded as an index fossil for lower Pliocene. The 154 unnamed marine unit is rich in globigerinids and open-marine benthic foraminifers, in contrast to the more restricted microfauna of the overlying Bouse Formation (Smith 1970, Olmsted et al. 1973). Winterer (1975) made paleobathymetric interpretions of microfaunas from the "older marine" unit in two wells, the CBA Federal-1 in the San Luis

Basin, and LCRP-29 in the Fortuna Basin. She characterized paleobathymetry of the CBA Federal-1 as "open marine shelf" (-200 m), and LCRP-29 as "at least 5000 m [sic]". The second figure may be a typographical error, as a paleodepth of >500 m seems more reasonable, in view of the

relatively shallow paleodepths of the Bouse and Imperial Formations.

Cuttings from this unnamed marine unit from the LCRP-29 well

(Olmsted et al. 1973) were examined by the author at the Arizona Bureau

of Geology. Sand grains in cuttings from this unit are predominantly

medium to very fine grained quartz, including translucent pink and

orange grains, and are typically subrounded. In contrast, sand from the

stratigraphically lower nonmarine sedimentary section in LCRP-23 and

LCRP-26 is angular and more feldspathic. These observations suggest

that sands of the Fortuna and San Luis Basins can be subdivided into

"C" and "L" suites, as in the FCV section.

In summary, the "unnamed marine unit" in the subsurface of the

Yuma Desert has the following characteristics in common with the Wind

Caves member: (1) a relatively well rounded, quartz-rich, "C"-suite

sand composition, associated with argillaceous beds; (2) an outer neritic to bathyal foraminiferal fauna and abundant planktonic foramin-

ifers; (3) presence of the planktonic foraminifer Sphaeroidinella 155 dehiscens, and (4) an overlying mud-rich marine unit with a more restricted, planktonic-poor microfauna (Mud Hills member in the FCV section; Bouse Formation in the Yuma Desert). For the same reasons the

"unnamed marine unit" is also quite different from the Latrania and

Lycium members. The angular discordance between the sand-rich and mud-rich units observed in the Yuma Desert (Olmsted et al. 1973) is not observed in the FCV section. However, this difference may be accounted for by local tectonic differences, related to different timing of nearby strike-slip faults, discussed in the last chapter.

Mud Hills member (Pm) (new name)

The Mud Hills member comprises two lithofacies: massive, monotonous, greenish gray (5 GY 6/1) to grayish olive (10 Y 5/2) claystone without apparent bedding or lamination, and yellowish-gray (5 Y 7/2), rhythmite-bedded clayey siltstone to silty, very fine sandstone (Plate

7). Although clean sandstones are scarce, the claystones and siltstones are very low in mica and similar to lithologies associated with "C"- suite sandstones in the overlying Deguynos member; the Mud Hills member is classified as a "C"-suite unit. In the FCV area the massive claystone facies is volumetrically dominant and characteristic of the lower part of the member, while the rhythmite-bedded siltstone facies is more typical of the upper part. These facies differences, however, are obvious only in fresh cut-banks of washes. Over most of the outcrop area this unit is deeply weathered and uniform in appearance, typically with a cracked, "popcorn-ball" surface; underlying lithologic differences are revealed only by subtle color variations (5 Y 7/1-3) on 156 weathered surfaces. Secondary vein-filling selenite is abundant in this unit, and weathered slopes are commonly littered with clear, reflective selenite plates.

This unit is named for Mud Hills Wash; it is also well exposed along Fish Creek Wash, North Fork, and Lycium Wash. It is equivalent to the "lower submember" of Woodard's (1963) Deguynos member, and the lower part of Stump's (1972) Burrobend member (Figure 1.9). It overlies the Stone Wash Member in the Lycium Wash Drainage; the two units interfinger to a minor extent. In the Mud Hills area it overlies the Wind

Caves member with a gradational contact. In the CIA the Mud Hills member overlies the Latrania member with sharp contact, except in the

Kamia Creek area, where it overlies the Kamia Creek member. In all cases, the base of the Mud Hills member is placed at the top of the highest sandstone or conglomerate bed at least 2 cm thick. The top of this member is placed at the base of the first resistant coquina or sandstone bed of the Deguynos or Lavender Canyon member.

Thickness of the Mud Hills member is difficult to measure directly because of the general lack of or poor exposure of bedding. Therefore, thicknesses were estimated primarily from outcrop width and the dip of bracketing units. On this basis, a maximum thickness of -600-700 m was estimated in the vicinity of Fish Creek Wash and

North Fork. The member thins to the northwest, and apparently pinches out between the Stone Wash and Deguynos members west of Coral Wash. To the southeast, it thins to -200 m south of Red Rock Canyon, to -50 m in 157 the Kamia Creek Section (KC-1, Figure 4.10), and is probably even thinner, though poorly exposed, along Barrett Canyon Wash.

Sedimentary facies

The massive claystone facies is characterized by lithologic uniformity (slightly silty claystone) and absence of lamination or bedding. Upward gradation into silty rhythmites is exposed along North

Fork. Well developed rhythmites are best exposed along Fish Creek Wash

(NF-6, Plate 7). The thickness of individual rhythmite cycles ranges

from 10 to 30 cm. In most rhythmite sequences the uniformity of cycle

thickness between adjacent cycles is striking; differential erosion

imparts a ribbed or corrugated appearance to fresh, vertical exposures.

Rhythmite sediments range in texture from silty claystone to clayey

siltstone, to sandy siltstone, to silty, very fine grained sandstone.

Textural gradation of individual rhythmite cycles is symmetric; they

grade upward from slightly finer to slightly coarser sediment, then

upward into finer sediment again. Rhythmite sequences are characterized

by the absence of sharp contacts and they are not classical graded beds

or turbidites. In some cycles, thinner (-1 cm or less) subcycles can be

recognized. Wisps of ripple lamination are locally apparent. Overall,

rhythmites appear to have been homogenized by bioturbation, presumably

by small organisms incapable of obscuring the primary cyclicity.

Discrete burrows, mostly small (<1 cm), are locally weathered out in uppermost fresh exposures along Fish Creek Wash and North Fork. In the 10 m and part of the Mud Hills member, upward-coarsening cycles between Chapter 6. 60 m thick can be observed (Plate 7); these are discussed in 158 Paleontology

Macrofossils are rare in the Mud Hills member, but aragonitic molluscs occur locally. Woodard (1974) reported "delicately preserved specimens of ... Nuculana, Tagelus, and Dentalium" from the "uppermost gray beds", probably of this unit. S. M. Kidwell (1986, written communication) reported a similar aragonitic fauna from gray claystones

exposed in the Carrizo Badlands along Andrade Canyon (Figure 4.2) and presumed to be equivalent to the Mud Hills member. This fauna included

Corbula, Dentalium, "Nuculana", Nucula, and Ensis. R. W. Norris (1984,

oral communication) discovered a sparse aragonitic assemblage in Fish

Creek Wash, consisting of small, unidentified pelecypods. Typical trace fossils in the rhythmite facies are vertical or near-vertical burrows,

<1 cm in diameter. One Teichichnus and a few "cone-in-cone" burrows were found in rhythmites of the Mud Hills member.

Planktonic foraminifers are less abundant in the Mud Hills member than in the Wind Caves member (Ingle 1974). Stump assigned the

Mud Hills (lower part of his Burrobend member) to foraminiferal Zone

N20 of Blow (1969), based on the appearance of Globorotalia multi- camerata and G. tosaensis. This age determination is apparently in conflict with magnetostratigraphic ages determined by Johnson et al.

(1983) (Chapter 6). Ingle (1974) interpreted the benthic microfauna as indicative of a maximum paleodepth of -50 m; Quinn and Cronin (1984) interpreted an "inner shelf-intertidal" paleoenvironment. 159 Lithologically equivalent units

Christensen's (1957) Butaca member of the Imperial Formation, described from the Coyote Mountains, was briefly examined in the field. It is lithologically indistinguishable from the Mud Hills member and occupies essentially the same stratigraphie position in the Coyote Mountains as the Mud Hills member does in the CIA (Figure 1.8). In the

San Felipe Hills and Yuha Basins, the base of the Deguynos member is apparently not exposed, but a Mud Hills equivalent may be present in the subsurface.

In the western Indio Hills, massive gray claystone, weathered to greenish-yellow, with secondary vein-filling selenite, and otherwise identical to the claystone facies of the Mud Hills member, crops out locally. These exposures have previously been referred to as Carrizo

Formation (Buwalda and Stanton 1930), Imperial Formation (Dibblee 1954,

Proctor 1968), and Willis Palms Formation (Stotts 1965). In fresh exposures this unit appears to be highly sheared, but, due to the absence of bedding, deformation is difficult to confirm. Sandstone is exposed locally near what may be the base of the section; these are fine grained, well sorted, subrounded, quartzose sandstones, and appear to be typical "C"-suite sandstones.

A sequence of marine mudstone encountered in wells of the

Fortuna and San Luis Basins has been assigned to the Bouse Formation

(Olmsted et al. 1973, Eberly and Stanley 1978, Winterer

1975). Foraminiferal faunas recovered from well cuttings of the Bouse in this area indicate upward-shoaling paleobathymetry, with outer 160 neritic in the lower Bouse above the outer neritic to bathyal "lower marine unit" (Winterer 1975). The Bouse grades up through a transitional unit into Colorado River alluvium (Olmsted et al. 1973). On the basis of these descriptions and examination of well cuttings by the

author, this unit appears more similar in lithology and stratigraphic position to the Mud Hills member of the FCV section than to the type

Bouse Formation (Metzger 1968, Metzger et al. 1973).

The "Imperial Formation" in the Sierra de los Cucapas report-

edly consists primarily of light to dark brown, massive marine clay-

stone with subordinate mudstone and siltstone and traces of bedding

(Barnard 1968). Beds of sandstone and coquina reportedly occur in this

unit as well; on the basis of Barnard's descriptions Mud Hills and

Deguynos equivalents cannot be separated.

Deguynos member (Pdg)

This unit consists of gray claystones and yellow-gray silt-

stones, similar to those of the Mud Hills member, gray bioturbated

mudstones, yellow-gray (5 Y 7/1-3) sandstones, and coquina beds. The

coquina beds, consisting primarily of Ostrea vespertina, Anomia

subcostata, and pectinids (Woodard 1963) in a sandstone matrix, are the

most conspicuous feature of this unit. They are more resistant to

erosion than most other facies of the "C"-suite sequence and form the prominent hogbacks (Plate 4), of which the Elephant Knees ridge is

most spectacular. Resistant coquina and sandstone beds typically

weather to dark brown (10 YR 4/2). 161 Hanna (1926) assigned similar coquina beds in the Yuha Basin to the "Yuha Reefs"; these were later incorporated into the Imperial

Formation (Woodring 1932, Dibblee 1954). Woodard (1963) named the Deguynos member after Deguynos Wash, and subdivided it into upper and lower submembers. The name is here restricted to Woodard's "upper submember"; the Mud Hills member having been substituted for his "lower submember" (Figure 1.9).

Lithofacies of the Deguynos member and bracketing units are intergradational, so that arbitrary boundary definitions are required for mapping. For the lower boundary, the first appearance of resistant, hogback-forming beds is the most useful mapping criterion. These beds consist of coquina or fine to medium grained, typically fossiliferous sandstone. To define the contact with the overlying Camels Head member, the beginning of reddish mudstones (at least 2 m thick) is a convenient mapping criterion. Although Woodard (1963) did not explicitly define his unit boundaries, the boundaries as defined here correspond closely with his mapped contacts.

The complete Deguynos member is well exposed only along Fish

Creek Wash (Figures 4.19-4.21) and North Fork (Figures 4.22-4.24), where the total thickness is approximately 300 m. The lower part of the section is also well exposed in the Elephant Knees area (Figure

4.25). To the northwest, on the southeastern flank of the Vallecito

Mountains, the Deguynos member probably interfingers with or grades laterally into the Stone Wash member and/or Camels Head member, but these relationships are obscured by faulting and alluvial-fan terraces 162

40 80 120

Pdg -35 -75 115 I -155

-30 70 110 -150 •

'11CC

-25 65 105 -145

-20 -60 100 •140

-15 FC-1- 127 95 -135

-10 -50 90 -130 1 Pdg --= Pm

-5 45 85 -125

C 120 0 40 80 FC -7

Figure 4.19. Graphic log of measured section FC-7, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plates 7 and 8. 163

40 IC 80 120

35 75 -115

30 70

25 65

FC-7-94.5

Pdg ' C o

FC-1 (part)

Figure 4.20. Graphic log of measured section FC-1, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plates 7 and 8. 164

20C1 240

:13

-195 B -235

-190

-185 265 Pch Pdg

-180 260

Coln ph

-175 255

-170 250

245 0 0 oh -165

Pdg C160 240 complex faulting here to FC-1-153 FC— 1 (part)

Figure 4.21. Graphic log of measured section FC-la, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 165

-15 55 -95

-10 rtlif vI stcsil_çj 50 _90 ss ss VS Pdg

-5 -125 45 z -85

-120 40 -80

-115 35 -75

11€- Gy,Th?,Te

- -10 30 -70 -110

- -15 25 -65 -105

rf,-4-48.5

--20 -100 20 -60

Pm NF-3

Figure 4.22. Graphic log of measured section NF-3, North Fork. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plates 7 and 8.

166

6 A .....,,_„_...,..... * U H II • II .*Thiph 1 -115 - --I.* - -155 60 ph fl 5 i%11 5 Gy•Tb7 ph

-110 -150

. — • g n°I u Y.

-105 145

large offset NF-4a 180

intervening thickness difficult to measure

175 for upward continuation 40 _ see Plate 7 offset /////

170

35 75

165

30 70

160

11- Thgc Pdg 25 65 NF-4

Figure 4.23. Graphic log of measured sections NF-4 and NF-4a, North Fork. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plates 7 and 8. 167

offset -55 • 95

sts 15115;5),j

—> NF-6-0 90 -50

offset Pch

85 -125 -45 large offs et

80 -120 -40

// /

75 -115 -35

Pch Pdg

"ifing-r— 110 It7e

65 -105 -25

60 -100 20

55 95 NF-5

Figure 4.24. Graphic log of measured section NF-5, North Fork. Duplication of numbering is due to re-measuring. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 168

c 40

80 120

-35

-75 -115

/ -30

- 70 -110

-25

-65 -105

of fset -20

-60 -100

m f vi sft...t.sj Sc ss ss -55 -95 135 Pdg

-50 -130

off set -5

-45 -85 -125

— ? C for underlying section see Plate 7 C 40 80 C 120 EK -2

Figure 4.25. Graphic log of measured section EK-2, Elephant Knees hogback. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 7. 169 (Figure 1.12, Plate 4). The Deguynos member is nearly continuously exposed into the Carrizo Impact area. The lower part of this member changes facies southeast of the Elephants Knees and becomes a distinct, mappable unit, which is referred to there as the Lavender Canyon member (Figure 1.8, Plate 4).

Sedimentary facies and paleocurrents

The lower part of the Deguynos member is characterized by

upward-coarsening cycles, typically 10 to 30 m thick, of claystone,

siltstone, and sandstone or coquina (Figures 4.19, 4.20, 4.22, 4.23,

4.25, Plates 7, 8). The uppermost part of the Mud Hills member also

contains such upward-coarsening cycles, with a maximum thickness >50 m, but without resistant sandstone or coquina beds, by definition. Upward-

coarsening cycles are sufficiently distinctive and laterally continuous

to permit physical correlation between closely spaced measured sections

(Plate 7). Some of these cycles can be correlated with confidence from

the eastern Elephant Knees to North Fork, a distance of nearly 5 km.

A typical upward-coarsening cycle consists of the following

lithofacies, although considerable variation on this basic theme is observed:

1. The lowest facies of most cycles is a massive, featureless, gray claystone similar to claystone of the Mud Hills member. In some cases, the contact with underlying sandstone is somewhat bioturbated, resulting in a decimeter-scale, upward-fining gradation. These claystone sequences are typically slightly more resistant to erosion than 170 are the siltstone sequences, and account for the "knees" on the northeast-facing slope of the Elephant Knees (Plate 7).

2. Massive claystone typically grades up into bioturbated mudstone, siltstone, and sandstone. In many exposures, hints of

"ribbing" are apparent in outcrop, suggesting rhythmite bedding, and in some exposures rhythmite bedding is obvious. In one such exposure, in

Loop Wash near its mouth (FC-1-6-12, Figure 4.20), highly concentrated small burrows are weathered out in relief in a rhythmite sequence.

Spectacular exposures of rhythmites in upward-coarsening cycles occur

along Lavender Canyon; individual cycles attain a maximum thickness of

-30 cm and maximum grain size of very fine grained sandstone, greater

than their counterparts in Fish Creek Wash. The thickness of rhythmite

cycles may systematically increase or decrease up-section within an

upward-coarsening sequence.

3. Many upward-coarsening cycles are topped with laterally

discontinuous, thick beds of sandstone or coquina. Characteristic

structures include trough cross-stratification, large-scale, low-angle

(10° or less) to high-angle (up to 25°) foresets, and mud flasers.

Where exposure permits, these beds appear to be lenticular, thinning or lateral even pinching out over distances of hundreds of meters. The

extent of typical resistant beds can be inferred by physical corre-

lation of upward-coarsening cycles (Plate 7). In this cross section,

the most continuous resistant unit forms the Elephant Knees hogback, this resistant and can be traced for nearly 3 km. On closer inspection, coquina unit actually consists of coalescing, lenticular sandstone and 171 some with bodies, sharp, scoured bases and others with gradational bases. Other resistant units are less continuous, commonly 1 km or less in lateral extent (Plate 7).

Coquina beds are composed of shell material in varying concentrations and states of preservation, from abraded fragments to un- abraded whole valves. In the FCV section, typical coquina beds consist exclusively of single valves and fragments, with a strong preferred orientation parallel to bedding, and a sand matrix. The sand matrix of many coquina beds is anomalously coarse and may contain admixed

"L"-suite sand, granules, and pebbles of granitoid or gneiss composi-

tion. S. M. Kidwell (oral communication) observed coquina beds in the

Coyote Mountains of almost pure shell material, including articulated

Ostrea shells. Coquina beds with articulated Anomia were observed by

the author on the southwest flank of the Coyote Mountains.

Basal contacts of these sandstone and coquina units show

considerable variability, even within a single bed. In some cases, the base is gradational with underlying rhythmite-bedded sandstone

(e. g. NF-3-0-17, Figure 4.23). More typically, the basal contact is

sharp but concordant. Locally, the basal contact is not only sharp but strongly discordant with underlying bedding. In other cases, the rhythmite sequence is absent and the sandstone or coquina bed overlies claystone directly. Sharp basal contacts are typically underlain directly by concentrations of sand-filled burrows, and bioturbation may blur otherwise sharp contacts. Upper contacts are rarely as well 172 exposed, and the upper surfaces of resistant beds typically form weathered dip slopes with dark brown desert varnish.

Paleocurrent directions were measured from cross-stratification in many of these sandstone and coquina beds (Figure 4.26). These data show a strong southerly mode, although a small secondary mode to the north is also observed. Medium-scale trough cross-stratification and

larger-scale (up to 6 in thick), lower-angle cross-stratification show similar orientations.

Higher in the section, the Deguynos member does not exhibit

such obvious large-scale organization (Figures 4.21, 4.23, 4.24). The

upper Deguynos member is characterized by gray claystone and relatively

thin beds of bioturbated siltstone and sandstone. Shell material,

similar in faunal composition to that of the lower Deguynos, occurs in

thin beds with a sandy matrix, but just as commonly in a silty or

clayey matrix, with random or weakly preferred orientation. Overall,

the amount of bioturbation is greater than in the lower Deguynos.

Lateral geometry of sedimentary bodies in the upper Deguynos is diffi-

cult to discern because of limitations of exposure.

High in the Deguynos, intervals dominated by very thin inter-

bedding of ripple-laminated sandstone and mudstone appear

(e. g. NF-48-151-158, Figure 4.23). Reineck and Wunderlich (1968)

referred to similar bedding types in modern tidal flats of the North

Sea as lenticular bedding, wavy bedding, and flaser bedding, depending

on the relative abundance of sand and mudstone. In this paper this association of bedding types will be referred to as wavy bedding (sensu 173

20%

n=67

ripple foresets

ripple crests

Camels Head member - Camels Head member - trough cross-stratification data from *wavy bedding" (s.l.) only - from thick sandstone (includes some data from bodies upper DeGuynos member)

20% 10%

n=150 n=8

ripples

Lavender Canyon member lower Deguynos member -

sole marks sandstone & coquina beds

Figure 4.26. Paleocurrent roses for Deguynos, Lavender Canyon, and Camels Head members. 174 lato). Paleocurrent directions from wavy bedding of the Deguynos member are discussed under the Camels Head member, where wavy bedding is a characteristic bedding type.

Paleontology

Previous studies of this unit have focused much attention on

the coquina beds; macrofossils have been listed by Smith (1962),

Woodard (1963), and Stump (1972). S. M. Kidwell (1986, written communi-

cation) provided a preliminary faunal list from the Elephant Knees and

North Fork areas (Table 4.2). In addition, Christensen (1957) and

Bell-Countryman (1984) have listed macrofossils from similar coquina

beds in the Coyote Mountains. The dominant species are Ostrea

vespertina, Anomia subcostata, Pecten deserti, and Pecten sp. (Woodard

1963). Christensen (1957) also reported Oliva spicata, Turritella

imperialis, Conus sp., and Ostrea heermanni from coquina beds in the

Coyote Mountains.

Most previous investigators have overlooked in situ molds and

casts of articulated pholads, which are common in mudstones directly

below sandstone beds. The shell material is not preserved and these

pelecypods rarely occur in coquina beds, which is probably why they in have been overlooked. Hanna (1926) listed the pholad Barnea costata strati the Imperial Formation, but did not give a specific locality or

graphic position. Woodard (1963) mentioned "burrowing pelecypods mew- embedded in situ in the calcareous siltstone" in the Camels Head

ber, which may refer to pholads. In virtually every case the pholads with are articulated and oriented perpendicular to bedding in mudstone, 175

Table 4.2. Tentative faunal lists for Deguynos member (Pdg), based on field identification by S. M. Kidwell (1986, written communication).

Lower North Fork (upper Deguynos member)

Ostrea vespertina Anomia subcostata Pecten spp. Balanus encrusting bryozoans on Pecten, whelk Polydora and Cliona borings in oyster Cerithium (rare) Crucibulum (molds) serpulid on Pecten Pinna? (prismatic shell material) colonia Astrangia (coral) attached to Pecten, Anomia, Pinna, and oyster fragments

Elephant Knees area (lower Deguynos member)

Ostrea vespertina (attached to oysters, Pecten, Anomia; with Cliona borings Pecten spp. (with Cliona and Polydora borings) Anomia subcostata (with attached oysters, Cliona and Polydora borings, rare bryozoans) Crucibulum molds Cerithium (rare) ?Barbatia (byssate bivalve) Corbula (small, interstitial dweller)

At thin, western edge of Elephant Knees coquina, add:

Encope Bal anus pho lad bone fragments coral (3-5 cm, unidentified)

Coquina bed at EK-2-102 (Figure 4.25)

Ostrea vespertina (worn/abraded, with heavy Cliona, Polydora borings) Ostrea heermanni with lithophagid borings barnacle clumps and attached to Pinna shark tooth, vertebrae, bone fragments echinoid spines (regular) Pecten (few) Turritella pholad mold 176 the anterior end downward. Multiple pholads can occur at the same horizon, with fairly uniform spacing. In an unusual occurrence at NFC-4a-156.5 (Figure 4.23), pholad casts occur in a network of

sand-filled Thalassinoides burrows. Some individuals underfit the burrows, suggesting that the original burrows were not created by the pholads.

The dominant trace fossils in the Deguynos are sand-filled

Thalassinoides and Gyrolithes burrows. Like the pholad casts, these

burrows typically occur in mudstone just below sandstone beds or

associated with wavy and lenticular bedding. Many of the unidentified

sand-filled burrows are probably Thalassinoides in cross-section, where

exposure is insufficient to show the diagnostic branching structure.

The Deguynos microfauna is similar to that of the Mud Hills

member. Stump (1972) assigned this part of his Burrobend member to

foraminiferal Zone N21 of Blow (1969), based on the appearance of

Globigerina calida and Globigerinoides sacculifera. As in the Mud Hills member, Stump's age assignment disagrees with magnetostratigraphic dates of Johnson et al. (1983) (Chapter 6). Quinn and Cronin (1984) interpreted the foraminiferal assemblages as indicative of intertidal to brackish conditions

Mitchell (1961) described bones of walrus from a coquina bed east of Painted Gorge in the eastern Coyote Mountains. He placed this locality within the Capote member of Christensen (1957), which is apparently equivalent to the Deguynos member. 177 Lithologically equivalent units

As previously mentioned, coquina beds similar those of the Deguynos in the FCV section occur in the Coyote Mountains (Capote member of Christensen 1957), the Yuha Basin (Yuha Reefs of Hanna 1926), and at Shell Reef and nearby localities in the San Felipe Hills

(Imperial Formation of Dibblee 1984, and Quinn and Cronin 1984). At

Yuha Buttes in the Yuha Basin, some coquina beds include local concentrations of the thick-shelled oyster Ostrea heermani, which do not occur in coquinas of the FCV section. In other respects the Yuha coquina beds and associated strata are essentially indistinguishable from those of the typical Deguynos.

The "Imperial Formation" described by Barnard (1968) from the

Sierra de los Cucapas has the essential characteristics of the Deguynos member. It consists of sequences of light to dark brown, massive claystones with subordinate mudstone and siltstone and traces of bedding. It is typically deeply weathered and eroded into badlands topography.The

Imperial also reportedly contains cuesta-forming beds and lenses of oyster-shell coquina and sandstone with abundant oyster and pectinid shells. Most sandstones are relatively mature, which Barnard (1968) attributed to a Colorado River source. Other reported fossils include

"small clams and monocotyledonous plant remains" (Barnard 1968).

In the Indio Hills, a sequence of mudstone, fossiliferous sandstone, and coquina is exposed in a small area just west of Willis

Palms, near the mouth of Thousand Palms Canyon. Coquina beds are dominated by Ostrea, but fragments of pectinids, turritellids, and vermetid 178

(?) tubes are also present. This is the type exposure of Stotts' (1965)

Willis Palms Formation. However, it is essentially identical to

Deguynos sediments in the FCV section. Stotts (1965) interpreted his

Willow Hole Formation, which contains Pleistocene vertebrates, to

underlie the Willis Palms. On this basis the Willis Palms was assigned

to the Pleistocene, and correlation with the Imperial was

discounted. However, given the striking lithologic similarity to the

Deguynos member, and the paleogeographic incompatibility of the fossiliferous Willis Palms with other Pleistocene units such as the lacus-

trine Brawley Formation (Dibblee 1954, 1984), it is more likely that

Stotts (1965) misinterpreted the Willow Hole - Willis Palms contact.

Powell (1986) assigned the same marine unit at Willis Palms to the

Burrobend member (nomenclature of Stump 1972) of the Imperial Forma-

tion. He suggested a correlation with the Burrobend member in the

Coyote Mountains, based on the occurrence of Anadara carrizoensis, and

interpreted the age to be younger than the Imperial Formation at Super

Creek (discussed under Latrania member). These interpretations by

Powell (1986) are consistent with those of the author.

In the Fortuna and San Luis Basins, there is no clear evidence

for a Deguynos equivalent. In particular, cuttings and electric logs

(Olmsted et al. 1973) do not suggest the presence of coquina beds.

Lavender Canyon member (Plc) (new name) Impact From southeast of the Elephants Knees into the Carrizo and Area, a sandstone unit occurs between the Mud Hills mudstones unit is typical Deguynos beds (Figure 1.6, Plate 4). This sandstone 179 evidently the lateral equivalent of the coalescing, lenticular sandstone and coquina beds that form the Elephants Knees hogback (Plate

7). However, this unit can be mapped separately and is here named the

Lavender Canyon member. Smith (1962) mentioned the presence of this member along Red Rock Canyon, but did not name or map it separately.

This facies apparently does not occur in or adjacent to the Coyote

Mountains (Christensen 1957) or other areas outside the Carrizo Impact

Area, so it is a rather local facies.

Due to logistical constraints, a thorough investigation of the

Lavender Canyon member was not possible, but one section (RR-1) was

measured along Red Rock Canyon (Figure 4.27). The boundaries of the

unit are somewhat gradational and are not strictly defined here. In

section RR-1 the unit is -80 m thick; adjacent to Lavender Canyon, a

thickness of -130 m was measured. A few salient points should be made:

1. Sandstones are a typical "C" suite and very uniform in grain

size, with a mean size near the fine/very fine boundary (3 phi).

2. Few sedimentary structures are visible, although minor

parallel lamination and convolute lamination were observed. Bedding

ranges from thin to massive, the latter being more characteristic. there 3. Ostrea shells occur as rare, local concentrations, but laterally are no true coquina beds, in contrast to the overlying and

adjacent Deguynos member in which coquinas are common. Claystone

intraclasts are also locally concentrated.

4. Locally, some sandstone beds are clearly lenticular. How- blocks in situ, ever, weathering and erosion have rounded off sandstone 180

40 110 150

-35 105 145

s -30 100 11.C.

sm. T h

-25 95 of f set

-20 4—Th 90

15 85

-10 80

Plc

Pm -5 75

70 RR-1

Figure 4.27. Graphic log of measured section RR-1, Red Rock Canyon. Location is shown in Plate 4. 181 so that lateral bedding geometry is difficult to perceive. It is likely that the Lavender Canyon member consists of laterally coalescing, lenticular sandstone bodies, similar to the Deguynos member at the

Elephants Knees.

5. The Lavender Canyon member is underlain by featureless claystone/mudstone of the Mud Hills member, and is overlain by upward- coarsening rhythmite-bedded cycles of the Deguynos member. The latter

are spectacularly exposed in Lavender Canyon.

6. Load casts and sole marks occur at the base of some sand-

stone beds. Paleocurrent measurements of flute casts and sole marks

show a southeasterly mode (Figure 4.26).

Camels Head member (Pch)

The vertical transition from marine Imperial Formation to

nonmarine Palm Spring Formation takes place over -300 m. Woodard named

this transitional unit the Camels Head member after Camels Head Wash,

where it is rather poorly exposed. Much better exposures are found in

Fish Creek Wash (Figures 4.28-4.30), North Fork (Figures 4.31-4.33) and

its tributaries Mollusk Wash and Jackson Fork. To the northwest this

unit interfingers with the Jackson Fork member in the Jackson Fork but drainage. It is locally well exposed in the Carrizo Impact Area,

its full extent there has not been mapped due to logistical member constraints, and it is left undifferentiated from the Deguynos

(Plate 4). this In the field the most obvious characteristic of

transitional unit is an up-section change in color. Claystone changes 182

Cg voli cm vf ss ss ss ss ss 40 80 120 Pch

-35 75 115

..... ,

-30 70 110

a:=D25 -25 65 105 '

-20 60 100

r oBe

-15 55 95 _ g_

50 90

......

/ / / /

-5 45 85 / / / /

Pch c-o 40 80 25.5 rn above NF-1-266 FC-2

Figure 4.28. Graphic log of measured section FC-2, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 183

C m f of srf,․)„csj ss ss so os 40 Pch d -115 155

«9 -75

-110 150

-30 -70

-105 145

-25 -65 E ==Ei -100 --f 140

• • ......

-20 60 z o „ . • , 0==, -95 135

-15 -55

-90 130

-10 -50

-85 125

-5 -45

-56 120

offset Pch o C 40 FC-3 (part)

Figure 4.29. Graphic log of lower part of measured section FC-3, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 184

-195 -275

-190

-185 265

FC —3a

-180 -260

R _175 -255 -295

offset(/'

-170 -250 -290

:8 ZB

285 °B -165 -245 = -13 = - Pd Pch

-240 280 z -160 -; — FC-3 (part)

Figure 4.30. Graphic log of upper part of measured section FC-3 and section FC-3a, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 185

10 ph

-5

Pch

NF-5a-128 4—> NF-6

Figure 4.31. Graphic log of measured section NF-6, North Fork. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 186

Cg VC SS SS 40

offset -35

-30

-25

-20

-15

-10

Pch Co

NF - 7 (part)

Figure 4.32. Graphic log of lower part of measured section 14F-7, North Fork. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 187

Cg VC SS SS Pd 120 160 -200 ...... ••••

Pch 115 -155 -195 NE-7 (part)

-150 -190

-145 -185

-140 -180

-175

Pd Pch ..... ••• -170 g -130

-125 -165

...

C 160 c 120

Figure 4.33. Graphic log of upper part of measured section NF-7, North Fork. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 188 from gray (5 Y 5-7/1-3) in the Deguynos member to reddish brown (5 YR 4-6/2-4) in the Diablo member; sandstone changes from yellowish gray (5 Y 7/2) to pale orange (5 YR 7-8/2-4). The Camels Head member contains both of the end-member claystone colors as well as intergradations, with the percentage of reddish claystone increasing up section. There-

fore, the base can be defined at the lowest occurrence of reddish

claystone or pale-orange sandstone at least 2 m thick (Figures 4.21,

4.24), and the top at the highest occurrence of gray claystone at least 2 in thick (Figures 4.30, 4.33).

Sedimentary facies

The Camels Head member has three dominant lithofacies which occur throughout the section.

1. Massive claystones range in color from gray to red, with red

increasing in percentage up-section.

2. Wavy-bedded (sensu lato) sandstone-siltstone-claystone

sequences are the most diagnostic sedimentological feature of the

Camels Head member. A continuum of bedding types is observed, including

laminated claystones with silt laminae, lenticular bedding with siltstone or sandstone lenses, wavy bedding sensu strictu, flaser bedding, and ripple-laminated sandstone without mud. Bioturbation in these sequences is generally minor. As in the massive claystone sequences, claystone color in the wavy-bedded sequences changes from predominantly gray (low in the unit) to predominantly reddish (high in the unit). Paleocurrent measurements from ripple-crest and ripple-foreset orientations (including data from wavy bedding in the upper Deguynos 189

member) show a dominant southerly mode, with a lesser northerly mode

(Figure 4.26). Individual wavy-bedded sequences typically show a strong

dominant direction; truly bimodal sequences were not observed, and

herringbone ripple lamination is rare.

3. Sandstone beds occur throughout the section, typically but

not exclusively in lenticular upward-fining cycles that average -5 m in

• thickness (Figures 4.21, 4.24, 4.28-4.33, Plate 8). The dominant struc-

tures in these sandstone units are parallel lamination, convolute

lamination, low-angle cross-stratification, trough cross-

stratification, ripple lamination, and flaser bedding. Coarse basal

lags are common in upward-fining cycles, including claystone intra-

clasts, shell material (Table 4.3), and granitoid and gneiss pebbles.In

some cases the shells are encrusted with bryozoans. Silicified logs are

common in these sandstones units. At the top of many upward-fining

cycles, sandstone grades up through siltstone into claystone; these

zones are typically biotrubated. Paleocurrent readings from trough

cross-stratification in such sandstone bodies indicate a dominant

southerly mode, without a significant secondary mode (Figure 4.26). In

the upper part of the Camels Head member, these sandstone bodies are

virtually indistinguishable to upward-fining sandstones in the over-

lying Diablo member. shells A minor but significant lithofacies consists of Ostrea or in a mudstone matrix. In such beds the shells are typically randomly Clumps of poorly oriented. Shells in some beds are bryozoan-encrusted. mixtures of attached Ostrea shells were found in some beds. Homogenized 190 sand, silt, and clay in some beds suggest the possibility of intensive bioturbation.

Paleontology

Quinn and Cronin (1984) placed this unit within a "brackish" biofacies on the basis of the foraminiferal assemblage, part of the

upward-shoaling and upward-freshening trend they show for the Imperial

Formation in the FCV section. Bioturbation and discrete burrows are

much less common than in the Deguynos member, which tends to indicate a

rigorous paleoenvironment and is consistent with Quinn and Cronin's

interpretation. One horizon bearing in situ pholads was found

(CT-2-89). Bioturbation is most characteristic of the upper portions of

upward-fining sandstone bodies, and possibly in Ostrea-bearing mudstone beds.

In view of the microfaunal and ichnologic evidence, the

invertebrate macrofauna is surprisingly diverse and abundant. Woodard

(1963) listed the following species from a single locality (his B-8685)

near Loop Wash: Ostrea vespertina, Ostrea heermanni, Aequipecten

deserti, Anomia subcostata, Atrina stepheni, Pecten sp., Turritella

im erialis, Solenastrea fairbanksi, Eusmilia carrizensis, "polychaete

worm tubes", and encrusting brozoans. In this study, the typical

Ostrea-Anomia-pectinid assemblage was observed to dominate the macrofauna, but various gastropods (especially Fasciolaria sp.),

Balanus sp., and sand dollars (Encope) were also found, typically with coarse lags in sandstone beds. S. M. Kidwell (1986, written communica-

tion) provided a preliminary list of the Camels Head fauna (Table 4.3). 191

Table 4.3. Tentative faunal list for Camels Head member (Pch), based on field identification by S. M. Kidwell (1986, written communication.

Upper North Fork (NF-6, Figure 4.31)

Ostrea vespertina (with encrusting bryozoans) Pecten sp. (with encrusting bryozoans) Anomia subcostata (with encrusting bryozoans) Astrangia col Solenastrea? pebbles with encrusting bryozoans and oysters unidentified gastropods with encrusting bryozoans serpulids on Ostrea vespertina Astrangia-bryozoan-serpulid balls to 5 cm diameter Balanus Ostrea heermanni fragments Crucibulum? circumrotary bryozoans with serpulids 192 Bryozoan encrustation is common, both on reworked shells (Anomia, Fasciolaria, Oliva, Ostrea) and on in situ Ostrea shells in mudstone matrix.

Woodard (1963) also reported teeth of Equus (cf. Plesippus) and an early Plesippus or late Pliohippus from near the upper contact of this member.

Lithologically equivalent units

A similar transitional unit, the Descanso member, was described by Christensen (1957) from the Coyote Mountains, and is probably an approximate equivalent. North of Shell Reef in the San Felipe Hills, a transitional unit is poorly exposed between the Imperial and Palm

Spring Formations of Dibblee (1984). In the Yuha Basin, reddish, wavy- bedded sandstone-mudstone lithofacies, characteristic of the Camels

Head member, is locally exposed in cut banks of washes. This area is structurally complex, and the full outcrop extent of this unit was not determined. In the Sierra de los Cucapas, Barnard (1968) reported some difficulty distinguishing between the Imperial and Palm Spring Forma- tions; this may indicate that a transitional facies like the Camels

Head is present. North of the Salton Sea, no equivalents of either the

Camels Head member or overlying Diablo member were observed.

A transitional member between the Bouse Formation and overlying

"older alluvium" has been described from wells of the Fortuna and San

Luis Basins (Olmsted et al. 1973) on the basis of lithology and microfauna. Cuttings from this interval in wells LCRP-26 and LCRP-29 consist of a mixture of "C"-suite and "L"-suite sand. This evidence 193 suggests thaithe "transition zone" of Olmsted et al. is equivalent in lithology and stratigraphic position to the Camels Head member.

Jackson Fork member (Pi) (new name)

The Jackson Fork member refers to a small but paleontologically significant exposure of marine to nonmarine "L"-suite sandstone and

conglomerate which interfingers with the Camels Head member in the

Jackson Fork drainage basin. The Jackson Fork member grades down-

section and toward the northwest (Figures 4.34-4.36), from mixed "L"

and "C" sandstones, to exclusively "L" sandstones, to conglomerates

that onlap plutonic basement in a high-angle nonconformity (buttress

wiconformity). The Jackson Fork member is faulted against the strati-

graphically lower Deguynos member, so that its stratigraphic relation-

ship with the Deguynos is not apparent. The Jackson Fork grades upward

into the lithologically similar but nonmarine 011a member. Woodard

(1963) included the Jackson Fork member in his Camels Head member. This unit is most notable for a very prolific shell bed, and several unusu-

ally large and well-preserved camel tracks, first recognized by Paul

Remeika (1984, oral communication). The best continuously exposed

section occurs along Jackson Fork (Figures 4.34-4.36). No comparable unit is known to be exposed elsewhere in the Salton Trough region.

Sedimentary facies

The Jackson Fork member consists primarily of "L"-suite sandstone and conglomerate, but due to interfingering with the Camels Head member, the measured sections (Figures 4.34-4.36) also include 194

40 67 0...... 0.1.„&z1.:0. -

.0 -..-021:17..0• • ea • •• 04o • CO

poj PODP • • 0 • oo OP? I •

Ct CN ON

r? 0.0 gifZ, ///// vq.a :n15'AIV 5fa":413 r

0 • o 0-0—, a° ,

*c2 oP:ogot

00,0 00 .7, 0 , ;:go 2.1,4tv,q1,9,0.8°,3 ?. 00 0 0' °

op°60

•"' a b 0 o ° o D 0 • ° vo ° /pcf:›404, 0.0s .0 0:?c, ; C 7:Pin 06: 0 rrl CT-1 CT—l a

Figure 4.34. Graphic log of measured sections CT-1 and CT-la, Jackson Fork. Location is shown in Figure 1.12 and Plate 4. 195

cg oci c m I vi srtetsj ss ss ss ss ss 80 120 Pch _L

75 /-115

e=r, ••=., flume ems Cs camel tracks in talus blocks from this 30 Interval 70 -110

65 =— -105 fr = /

-100

Zarc.4-9,0427/,

. - :TO ..14;rea7 -95

././ 'Ave -90

minor Intervening faults 5 -85

rf-////Zaal-42:, C 80

Figure 4.35. Graphic log of lower part of measured section CT-2, Jackson Fork. Location_is shown in Figure 1.12 and Plate 4. 196

160 200 /MC

-155 195

,4%*5.1Mplf&WV

...... „ 190

...... ^ •

-145 185

11-II-11-1t— ph C

-140 180 220

-135 175 215

11, —130 170 210 -6

165 205

o ,-- PO, 200 T 120 160 CT-2 (part)

Figure 4.36. Graphic log of upper part of measured section CT-2, Jackson Fork. Location is shown in Figure 1.12 and Plate 4. 197 intervals of "C"-suite sandstone, siltstone, and associated reddish claystones. The base of the measured section consists primarily of polymictic boulder to pebble conglomerate. Boulders and pebbles are mostly subrounded to subangular, and consist primarily of granitoid plutonic rocks similar to nearby quartz diorite basement, but also include gneiss clasts. Matrix is a coarse to very coarse, poorly sorted, feldspathic sandstone. Parallel lamination and low-angle cross-stratification are visible in places in the sandstone matrix.

Above CT-1-20, marine indicators appear. The most significant of these

is the fossil bed at CT-1-24 (Figure 4.34), which forms a well exposed

dip slope. Slightly higher in the section (up to CT-1-56), several

Ostrea-cerithiid-bearing, pebbly sandstones occur. Some sandstone beds

in this interval are color-mottled and weather to a "hackly" surface,

suggestive of bioturbation. A thick sandstone bed at CT-la-12-19

(Figure 4.34) shows large scale accretionary cross-stratification

(epsilon cross-stratification of Allen 1963). This unit contains

dispersed Ostrea shells in conglomeratic sandstones.

Higher in the section, to CT-2-79 (Figure 4.35), the Jackson

Fork member consists primarily of thin to medium bedded "L" type

sandstones, with great variability of grain size. This sequence is

characterized by parallel lamination to low-angle cross-

stratification, some ripple laminations, and numerous small scour- Burrows and-fill structures, but little trough cross-stratification.

occur at some horizons, most notably an Ophiomorpha nodosa-burrowed

sandstone at CT-2-32.5-34.5, but there is little skeletal material. 198

Mud cracks and calcareous nodules (paleocaliche?) appear locally. A few debris-flow beds also occur in this sequence, characterized by poorly

sorted, angular sand, granules, and pebbles in a dark-olive clay-rich matrix.

Paleontology

A diverse invertebrate (mostly molluscan) fauna (Table 4.4) was

collected by S. M. Kidwell and the author from a single bed in the

Jackson Fork member (CT-1-23-24, Figure 4.34). Turritella and Encope

are particularly abundant in this bed, an unusual occurrence this high

in the Imperial Formation. Beds bearing cerithiid gastropods (Figure

4.34) are apparently unique in the Imperial Formation. Woodard (1963)

discussed this member under his Camels Head member, and reported Ostrea

vespertina, Anomia subcostata, Aequipecten deserti, Pecten sp., Balanus

cf. B. tintinnabulum, Murex sp., Pododesmus macroschisma, and ?Hanetia

sp. from localities in the Jackson Fork outcrop area.

Several large cloven hoofprints, probably of Camelops, were

observed on the underside of "L" sandstone beds, in some cases on talus

blocks. Some of these are of very high relief, whereas others are quite

subtle. Many of the observed camel tracks may be underprints, because

of the lack of clear definition. The tracks in this locality are larger

(-20 cm) than camel tracks previously known from "L" sandstones in the

011a member exposed at Camel Ridge (P. Remeika, 1984, oral communi-

cation). Bones of Came lops have been found in the Loop Wash area at

about the same stratigraphic level (Woodard 1963). 199 Table 4.4. Tentative faunal lists for Jackson Fork member (Pj), based on field identification by S. M. Kidwell (1986, written communication).

CT-1-24-25 (Figure 4.34)

Turritella (dominant taxon) Encope Dosinia Glycymeris Crassatella "Thais" in clumps Conus Oliva Olivella knobbed whelk (?Strombus) cowrie Evola (rare) naticid (no Cerithium, oysters, or barnacles)

CT-1-23 (Figure 4.34)

Ostrea vespertina (on cobbles and in sandstone) Cerithium (in sandstone) naticids (in sandstone) Anomia holdfasts on cobbles Crucibulum associated with gravel lags

(Note: little evidence of post-mortem mixing of faunas of hard and soft (sandstone) substrates; suggests low-energy conditions. CHAPTER 5

STRATIGRAPHY OF THE UPPER NONMARINE UNITS

The upper nonmarine units comprise the Palm Spring Formation, the Canebrake and Ocotillo Conglomerates and the Borrego and Brawley

Formations (Dibblee 1954). Woodard (1963) subdivided the Palm Spring

Formation into the informal Diablo, Tapiado, Hueso, and Vallecito members; his nomenclature will be followed here, with modifications

(Figure 1.9). As with the marine units, sedimentary provenance provides a convenient basis for subdividing the upper nonmarine section (Figures

2.9, 2.10): "L"-suite units (Canebrake and Ocotillo Conglomerates,

Tapiado, Hueso, and Vallecito members), and predominantly "C"-suite

units (Diablo member and Borrego and Brawley Formations). The 011a member (new name) is an extensive alternating-provenance unit that is

laterally transitional between the Diablo member and the Canebrake

Conglomerate or Hueso member.

Alternatively, these units can be subdivided into predominantly

lacustrine units (Tapiado member; Borrego and Brawley Formations) and

predominantly fluvial and alluvial-fan units (Diablo, 011a, Hueso, and

Vallecito members; Canebrake and Ocotillo Conglomerates). The Borrego

and Brawley Formations and Ocotillo Conglomerate do not occur in the

FCV area, but are extensively exposed to the north in the Borrego

Badlands, San Felipe Hills, and Superstition Hills (Plate 3). Of the

200 201 upper nonmarine units, only the Diablo and 011a were studied in detail, as these are critical for interpreting paleogeography of the ancestral Colorado delta. Description of the other units is based primarily on Woodard (1963) and Dibblee (1984), supplemented by field reconnaissance. Physical stratigraphic relationships among the upper nonmarine units are summarized in Figures 1.6, 5.1, 5.2, and Plate 3.

Diablo member (Pd)

The Diablo member is a very distinctive and extensively exposed unit of friable, pale orange (5 YR 7-8/2-4), fine-grained "C"-suite sandstones and siltstones, and reddish brown (5 YR 5/2-4) claystones.

In Woodard's (1963) definition, the Diablo includes a facies which consists of alternating "C" and "L"-suite strata exposed in the upper

Fish Creek Wash and Arroyo Tapiado drainages. He designated this mixed-provenance unit as the "upper submember" of the Diablo, and the rest as the "lower submember". These designations are somewhat misleading, as the alternating-provenance facies is more a lateral equivalent to the "C"-suite facies than an overlying unit. In the present study "Diablo" is restricted to Woodard's (1963) "upper submember", and

"011a member" is substituted for his "lower submember" (Figure 1.9).

The Diablo member grades downward into the transitional-marine,

"C"-suite Camels Head member and laterally into the alternating- provenance 011a member. The upper contact of the Diablo member in the

FCV section is sharp and laterally persistent, defined by the uppermost occurrence of pale orange "C"-suite sandstone, associated pale-orange siltstone, or reddish claystone. Overlying units, the Hueso and Tapiado 202

NORTH PALE0- NORTH A ALL PALE0- PALE0- CHANNEL CURRENTS VECTOR MEANS (N=623) (N=86) N=101

:Oal

N=38 Qa1 alluvium

0 Hueso POh member Tapiado Pt member PQc Canebrake PQh Conglomerate N=125 01 la Po member Diablo Pd member Camels Head Pch member Mud Hills & Pmd Deguynos mbrs. Wind Caves Pw members mpn., pre-deltaic 0 marine 5 pre-deltaic Mn nonmarine

Figure 5.1. Simplified geologic map of FCV section, showing distribution of upper nonmarine units and paleocurrent roses for "C"-suite all sandstone bodies in Diablo and 011a members. Paleocurrent data are from trough cross-stratification. "Paleo-north" arrow is based on paleomagnetic declinations from Johnson et al. (1983). 203

SAN FELIPE HILLS BORREGO BADLANDS Qb (lacustrine) Qo

Pb _ Pb (lacustrine) (lacustrine) • —

thick thick tongues transition zone transition zone

Pd Pd (delta plain) (delta plain)

VALLECITO BASIN

0.9 Ma PQh

pt/(lacustrine) 2.8. Ma APP.% 4•411%... Ard•n•nn 1.-suite AMMUMWnn _ Pd tongues (delta plain) 7./. II.Or4...... —_

AFAIAMdrAr...--

Figure 5.2. Schematic cross sections of upper nonmarine units in western Salton Trough (Plate 3). Diagonal ruling indicates Predominantly "L"-suite sediments. Note overall similarity of facies relationships between Vallecito Basin and Borrego Badlands, despite differences in nomenclature and lack of control on temporal correlation. 204 members, are olive, greenish, or gray, and consist of typical "L"-suite arkosic sandstones and biotite-rich siltstones and claystones. In the CIA the upper contact of the Diablo is not exposed, and the overlying units have probably been removed by erosion.

The Diablo member is the most extensively exposed Neogene unit in the Salton Trough (Plate 3); in much of this area it is essentially

synonymous with "Palm Spring Formation" as previously mapped (Dibblee

1954, 1984, Christensen 1957, Smith 1962, Hoover 1965). It is quite

homogeneous throughout its extent. In the FCV section the lower Diablo

is best exposed along Fish Creek Wash between Camels Head Wash and

Layer Cake Wash (Figures 5.3-5.5), and the upper Diablo is best exposed

along Arroyo Seco del Diablo in Diablo Canyon (Figures 5.6-5.9) and

locally along Arroyo Tapiado (Figure 5.10). This outcrop area extends east into the CIA, and the Diablo is extensively exposed along Deguynos

Wash. South of Carrizo Wash, good exposures can be found in the Carrizo

Badlands east of Canyon Sin Nombre (Plate 4). Numerous poorer exposures

are distributed around the Coyote Mountains (Christensen 1957). The

Diablo also occurs in the Borrego Badlands-San Felipe Hills region

(Palm Spring Formation of Dibblee 1984); it is particularly well exposed along Arroyo Salada and Tule Wash west of State Highway 86 (Plate 3).

Sedimentary facies

The Diablo member exhibits a considerable diversity of sedimentary structures and bedding types (Figures 5.3-5.10), but these can be summarized in terms of two major lithofacies: (1) thick-bedded to 205

V CI C m I I vf strtSI SS SS ss $s ss Pd or•-• 40 8 120 -160

35 -75 -115 -155 .13

040 30 -70 -110 -150

'13 -25 -65 -105 -145

20 -140

-15 -135

-130

-125 5 45 -

C 120 40

FC-4

Figure 5.3. Graphic log of measured section FC-4, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 206

PL -55

72,7p.vvaa

...

. .

-45

40 FC -4a

Figure 5.4. Graphic log of measured section FC-4a, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 207

FC-4b

Figure 6.6. Graphic log of measured section FC-4b, Fish Creek Wash. Location is shown in Figure 1.12 and Plate 4. This log is also shown in Plate 8. 208

DC-1 DC-2 DC-3

Figure 5.6. Graphic logs of measured sections DC-1 to DC-4, Diablo Canyon. Locations are shown in Plate 4. 209

DC-6

Figure 5.7. Graphic logs of measured sections DC-5 to DC-7a, Diablo Canyon. Locations are shown in Plate 4. 210

160

155

150

145

140

135

130

125

120

Figure 5.8. Graphic log of lower part of measured section DC-7b, Diablo Canyon. Location is shown in Plate 4. 211

200

• 8 CN 195

• CN

190

185

180

175

(minor hatIRMO 170

245 165

240 Pd 160 DC-7b (part)

Figure 5.9. Graphic log of upper part of measured section DC-7b, Diablo Canyon. Location is shown in Plate 4. 212

cg lvc c hi f I vi 'Tcsi vs ss ss ss ss

Pt N Pd 35

Pd :1 ...... AT-1

Figure 5.10. Graphic log of measured section AT-1, Arroyo Tapiado. Location is shown in Plate 4. 213 massive, fine to very fine sandstones, commonly in upward-fining cycles, and (2) sequences of reddish claystone, commonly interbedded with siltstone and/or sandstone. The Diablo does show some systematic up-section lithofacies variation, but not nearly as much as the underlying Deguynos and Camels Head members.

Massive to thick-bedded sandstone bodies in the Diablo exhibit upward-fining to blocky (sharp base and top) textural profiles. Upward- fining cycles are typically 5 to 10 m thick, although thinner cycles, many representing partial, truncated cycles, are common as well. The mean grain size is mostly lower fine to upper very fine. Basal portions of many cycles are significantly coarser grained, in many cases with admixed "L" sands, and the upper portions of many cycles grade up through siltstone into claystone.

Bases of sandstone bodies are in many cases defined by concentrations of intraclasts of red claystone and lesser calcareous nodules, admixed coarse "L" sand, pebbles, or rarely, bone fragments or shell material. Basal contacts are invariably sharp, and commonly show angular discordance with underlying strata. As much as 5 m of erosional relief (DC-6-26-31, Figure 5.7) was observed on basal contacts. Load casts occur at some basal contacts. Individual sandstone bodies vary considerably in thickness due to their lenticular geometry. Coalescing lenticular sandstone bodies are common, resulting in large aggregate sandstone thicknesses, but the basal contacts of individual cycles or lenticular bodies are generally well defined. In a few localities, suggestions of large-scale accretionary stratification (epsilon 214 cross-stratification) were observed (e. g. DC-4, Figure 5.6;

DC-7a-35-40, Figure 5.7), but this is uncommon.

Parallel lamination, convolute lamination, and low-angle

cross-stratification are the most common sedimentary structures in the

massive sandstone bodies. Ripple lamination is common in very fine

sandstones and siltstones in the upper portions of upward-fining

cycles. Many sandstone bodies, however, appear largely structureless

even on fresh surfaces, presumably due to a lack of textural or

mineralogic variation to accentuate laminations. Low-angle truncations

are common within sequences of parallel lamination, indicating non-

horizontal paleosurfaces. Heavy-mineral concentrations are most

typically associated with parallel lamination high within individual

cycles. Parting lineations are also commonly associated with parallel

lamination. Low-angle trough cross-stratification, with subtly concave-

upward stratification, is most typical of basal parts of cycles,

commonly in association with claystone intraclasts.

Convoluted structures up to 2 m in thickness are abundant in

Diablo sandstone bodies, in many cases obscuring primary sedimentary

structures. Convolute structures may occur at multiple levels in a

single upward-fining cycle. Some smaller-scale convolute structures

apparently represent oversteepened or fluidized foresets in trough

cross-stratification. However, in many cases, convolute lamination

apparently represents deformed parallel lamination. the structures Trough cross-stratification is less common than the bases of cycles. It previously discussed, and generally occurs near 215 is more common in coarser-than-average sandstones, as where "L" sands are admixed.

Silicified logs and smaller plant fragments are common in sandstone bodies of the Diablo member, as in the Camels Head member. Unlike the coarse material of basal lags, silicified wood is not generally confined to the lowermost parts of cycles. Fossil logs and other plant fragments are surrounded by rusty yellow to brown (10 YR 5-6/6) discoloration haloes, which makes them easy to locate.

A conspicuous feature of Diablo sandstone bodies is the

abundance of concretions which weather out in myriad shapes, including

spheroids, rods, discs, and ameboids. Concretions are locally concen-

trated in lag gravels such as the "Pumpkin Patch" in the San Felipe

Hills. These concretions result from local concentrations of carbonate

cement and may reflect pre-existing geochemical anomalies; however,

they do not appear to be related to any primary depositional feature of the sandstones.

Fine-grained sequences of claystone with interbedded siltstone

and very fine sandstone make up -20% of the Diablo member. Wavy bedding

occurs in the lowermost part of the Diablo member, just above the

Camels Head contact (Figures 4.30, 4.33), but is otherwise absent from

the Diablo member. Claystones are predominantly massive and reddish,

but thin greenish-gray to bluish-gray horizons also occur. Mud cracks

are locally common, defined by crack-filling sand or silt, resulting in

a brick-like appearance in cross section. Concentrations of calcareous nodules occur in some beds and horizons, and reworked septarian 216 calcareous nodules are locally common in intraclast conglomerates in basal lags of upward-fining sandstone bodies. Mud cracks, calcareous nodules and horizons, and calcareous intraclasts are most common in the upper part of the Diablo member than in the lower part, and are best seen in Diablo Canyon (Figures 5.7-5.10).

Interbedded siltstones and sandstones are in some cases laterally continuous and of uniform thickness; in other cases they are

quite lenticular. Small-scale ripple lamination is the dominant sedimentary structure, including climbing-ripple (ripple-drift) lamina-

tion. Larger-scale cross-stratification is rare in these sequences.

Beds and sequences of olive to gray, biotite-rich sandstones,

siltstones, and claystones, with the characteristics of typical

"L"-suite sediments, occur within the Diablo member, particularly along

Fish Creek Wash north of the Diablo Turnoff (e.g. FC-4b, Figure 5.5).

Sedimentary structures are predominantly large-scale tangential cross-

stratification and small-scale ripple lamination. These beds and

sequences are distal tongues of "L"-suite sequences of the 011a member.

The Diablo member is quite similar to the upper part of the

Camels Head member, into which it grades. However, the Diablo lacks the distinctive wavy-bedded lithofacies and gray claystone sequences that characterize the Camels Head member. Shell material in basal lags is much less common in the Diablo member than in the Camels Head member.

Paleocurrents

Despite the extensive exposure of the Diablo member, paleo- cluTent determination is not a trivial matter, due to the paucity of 217 trough cross-stratification, the tendency for weathering to obscure sedimentary structures, and the broad scatter of paleocurrent directions, which necessitates a large number of measurements to obtain statistically significant modal directions. Paleocurrent measurements were made from trough cross-stratification in upward-fining sandstone cycles; parting lineations and ripple lamination were also measured.

For individual sandstone bodies with sufficient data, vector means were computed. Data are divided into three areas (Figure 5.1): Area 1, in widely scattered localities along Fish Creek Wash (FC-3, 4, 4a, 4b),

Mollusk Wash (MW-1, 2), and Blackwood Basin Wash (BW-m localities), which collectively represent the lower part of the Diablo member; Area

6 along lower Diablo Canyon (DC-1 to 7b), representing the upper part of the Diablo member; and Area 4 along Arroyo Tapiado and adjacent small canyons (AT-1, AT-m localities), also representing the upper Diablo.

Paleocurrent directions from the lower Diablo (Fish Creek and

Blackwood Basin) show a high dispersion, but with a strong southerly mode (Figure 5.1). Paleocurrent directions for the upper Diablo (Diablo

Canyon and Arroyo Tapiado) show more southeasterly modes, again with large dispersions (Figure 5.1). Paleomagnetic-declination data from the

FCV section indicate -35° of post-depositional counterclockwise rotation (Johnson et al. 1983), as indicated by the "paleo-north" arrow in

Figure 5.1. Correcting for tectonic rotation would rotate the mean directions more toward the east. Composite results for "C"-suite sandstones will be discussed under the 011a member. 218 Paleontology

Vertebrate faunas of the Palm Spring Formation have been summarized by Woodard (1963) and Downs and White (1968). It is not always apparent from these discussions which observations pertain to which member. In particular, it is sometimes unclear whether bones were recovered from "C"-suite or "L"-suite units. However, the majority of collecting localities reported by Woodard (1963) are from the Hueso member, which may indicate a preferential preservation of bones in "L"-suite strata.

Woodard (1963) described the Diablo fauna as including molar

fragments of Equus cf. Plesippus and horse and camel bones, plus abundant chelonian (turtle) fragments and fish spines; he interpreted this fauna as Blancan in age. This is a rather sparse fauna compared to that reported from the Hueso member. Downs and White (1968) assigned a

Blancan age to this part of the Palm Spring Formation. Woodard (1963) also observed occurrences of Ostrea vespertina, Anomia deserti, and

Pecten deserti in the Diablo member. He characterized these fossils as

"stunted and abnormally shaped" (p. 94) and interpreted them as evidence of a "brackish water or moderately saline environment" within the

Diablo. In this study, these species were observed in the Diablo member only in basal lag deposits, typically in association with other coarse material including claystone intraclasts, coarse "L"-suite sand, gravel, and bone fragments. The shells are commonly abraded, but not significantly stunted or otherwise abnormal when compared to individuals of the same species in the Imperial Formation. Near the top of 219 the Diablo member in Arroyo Tapiado (Woodard's locality B-8690), shell material is unusually abundant and mixed with a variety of other gravel-size material. At this locality a clast of Ostrea coquina was found in the same bed, very similar to coquinas of the Deguynos member. From these lines of evidence, marine fossils in the Diablo member are thought to be reworked from the Imperial Formation. Silicified logs have been identified by Remeika, Fischbein, and Fischbein (in press) as the laurel Umbellelaria salicifolia

(Lesquereux) Axelrod, two species of Salicaceae (cottonwood, Populus

sp., and willow, Salix sp.), and the walnut Juglans pseudomorpha

Axelrod. Cottonwood, willow, and walnut were the dominant trees along

banks and levees of the lower Colorado River in its natural state

(Kniffen 1932). Laurel is more characteristic of upland chaparral in

California (Owen Davis, 1986, oral communication).

Lithologically equivalent units

Because the Diablo member is very distinctive in appearance, it

is readily identified in adjacent areas. The "Palm Spring Formation" mapped by Christensen (1957) to the north, south, and east of the

Coyote Mountains, and by Smith (1962) in the Carrizo Impact Area,

consists exclusively of Diablo-like lithologies with similar sequences

and structures. Outcrops of the Palm Spring Formation mapped by Dibblee

(1954) in the Yuha Basin also appear to be the Diablo member.

The Palm Spring Formation, as mapped by Dibblee (1984) in the

San Felipe Hills and Borrego Badlands, consists primarily of the Diablo member, but near the south and east flanks of the Santa Rosa Mountains 220 it consists primarily of "L"-suite sandstones. The latter deposits are equivalent in appearance to the 011a member of the FCV section. Hoover

(1965) mapped this facies near the Santa Rosa Mountains as "Facies B" of the Canebrake Conglomerate, rather than Palm Spring. His Palm Spring

Formation is equivalent to the Diablo member of this study. In the western Borrego Badlands and in Borrego Mountain Canyon (Borrego Moun- tain 7.5' Quadrangle), the Palm Spring Formation as mapped by Dibblee (1984) contains "L"-suite sandstones interfingering with "C"-suite sequences, and is mapped here as 011a member (Plate 3).

The "Palm Spring Formation" described by Barnard (1968) from the Sierra de los Cucapas is similar to the Diablo member of this study. Barnard reported that the Palm Spring consists of great thicknesses (locally >1500 m) of light gray to light brown, fine to medium grained, poorly indurated, cross-stratified sandstone (feldspathic litharenite) with siltstone and mudstone. The Palm Spring overlies the marine Imperial Formation or, locally, igneous and metamorphic basement; it grades upward into the "Canebrake Formation". Outcrops of the

Palm Spring are typically weathered and eroded to low-relief, hummocky topography. As in the Diablo member, silicified wood and concretions are locally abundant in the sandstone; concretions occur in a variety of shapes and are locally concentrated in lag gravels or "pumpkin patches". Unlike the Diablo member, however, sandstones of Barnard's

(1968) Palm Spring Formation reportedly contain rounded pebbles (up to

15 cm diameter; most <2 cm) of chert, quartzite, and other extra-local lithologies. These include fossiliferous Paleozoic limestones with 221 crinoid columnals, bryozoans, brachiopods, and corals; lithologies are similar to the Kaibab and Redwall Limestones of the Colorado Plateau. Cross-stratification suggests a southerly to westerly transport direction (Barnard 1968).

"Palm Spring" has also been applied to nonmarine terrigenous sediment overlying the Imperial Formation in the Indio Hills (Dibblee

1954, Proctor 1968). However, these sediments are mostly coarse- grained, conglomeratic, and completely unlike the Diablo member. On the basis of field reconnaissance and published descriptions, Diablo equivalents are believed to be completely absent from the Indio Hills and San Gorgonio Pass regions.

Olmsted et al. (1973) described a nonmarine sequence of "older alluvium" overlying their "transitional unit" in the Fortuna and San

Luis Basins; this unit consists at least in part of Colorado River- derived alluvium and is therefore stratigraphically analogous to the Diablo member.

011a member (Po) (new member)

The 011a member consists of alternating and interfingering sequences of "C" and "L" suite sandstones, siltstones, and claystones

(Figures 5.11-5.14). Contrasting coloration, mineralogy, and texture of the two sedimentary suites are readily apparent in the field and make differentiation of the two provenances straightforward. Woodard (1963) named this unit the "upper submember" of his Diablo member, which is somewhat misleading, as previously discussed. It is regarded here as sufficiently different from Woodard's "lower submember" and 222

Figure 5.11. Graphic log of lower part of measured sections LC-1, Layer Cake Wash. Location is shown in Figure 1.12 and Plate 4. 223

240

200

-235

195

-230 VC f sitsj,c_si SS SS SS ssss Po 190 265

-225 r--

185 260

-220

-180 255

-215

-175 250

large Oftm

-170 -210 245

240 -165 -205

Po 235 offtet; 160 200 LC— 1 (part)

Figure 5.12. Graphic log of upper part of measured section LC-1, Layer Cake Wash. Location is shown in Figure 1.12 and Plate 4. 224

OW-1 OW-5

Figure 5.13. Graphic log of measured sections 0W-1 to 0W-5, 011a Wash. Locations are shown in Plate 4. These sections were selected to show "C"-suite sandstone bodies bracketed by "L"-suite strata. 225

OW-6 OW-7 OW-8

Figure 5.14. Graphic log of measured sections 0W-6 to 0W-8, 011a Wash. Locations are shown in Plate 4. 226 sufficiently important for regional paleogeography (Chapters 6 and 7) to deserve a separate name.

Outcrops of the 011a member are somewhat more resistant to erosion than those of the Diablo and are characterized by higher-relief topography. Consequently, the 011a is well exposed in a number of areas in the FCV section, including upper Fish Creek Wash, Layer Cake Wash (Figure 5.11, 5.12), 011a Wash (Figures 5.13, 5.14), Sandstone Canyon, Syncline Wash, Arroyo Tapiado (trunk and north fork), and upper Arroyo

Seco del Diablo. South of Carrizo Wash, the 011a is well exposed (and easily accessible) along Canyon Sin Nombre, and also crops out just south of Sweeney Pass.

The lateral boundary between the 011a and Diablo members is

intergradational and difficult to define precisely; the 011a is

arbitrarily defined here by the presence of at least 20% "L"-suite

sandstone and siltstone in the sequence. The lateral boundary with the

Canebrake is also broadly gradational; it is defined here as the change

from sandstone to conglomerate as the dominant lithology. The 011a

grades downward into the Jackson Fork member which it closely resem-

bles; the contact is placed by projecting the Diablo-Camels Head contact to the northwest. The upper contact of the 011a in the FCV section, in contrast to the other contacts, is sharp and unequivocally defined by the uppermost occurrence of "C"-suite sandstone beds which crop out in the upper Arroyo Tapiado drainage (e. g. AT-m-4). These

Upper "C"-suite beds are armored with abundant concretions and form fairly prominent hogback ridges which can be easily mapped on aerial 227 photographs. The top of this "C" sequence is taken as the top of the 011a member. In Canyon Sin Nombre, however, the the upper contact of the 011a as defined by the uppermost occurrence of "C"-suite sandstones in an interfingering contact with the Hueso member (Plate 4).

Sedimentary facies and paleocurrents

"C"-suite sequences in the 011a member are very similar in

appearance to sequences in the Diablo member, and most observations about the Diablo also apply to these sequences in the 011a. The prin-

cipal difference is that basal portions of upward-fining sandstone

cycles in the 011a typically contain a higher percentage of admixed "L"

sand and gravel than is characteristic of the Diablo. These coarse

basal-lag concentrations can usually be recognized as part of upward-

fining "C"-sandstone cycles by (1) the presence of red claystone

intraclasts, and (2) upward gradation into finer grained, pale orange

sandstone. Trough cross-stratification is common in these basal-lag

concentrations, which facilitates paleocurrent measurement.

Paleocurrent data were obtained from "C"-suite upward-fining

sandstone cycles in two areas (Figure 5.1): Area 2, including Layer

Cake Wash and 011a Wash, and Area 3, Arroyo Tapiado. Data from the

first two areas were combined to represent the lower portion of the

011a member; Arroyo Tapiado represents the upper portion of the 011a.

These data show a dominant southerly direction for both high and low

Parts of the member, similar to results obtained for the Diablo. For

sandstone cycles of the Diablo and 011a members with sufficient data, 228 vector means were calculated (fig. 5.1); again, a southerly mean direction is indicated.

"L"-suite sequences in the 011a member (Figures 5.11, 5.12) are broadly similar to those in the Jackson Fork member. These consist of very light gray (N 7-8) to olive gray (5 Y 6-7/1) sandstones, biotite- rich siltstones of similar color, and darker olive (5 Y 4-7/1-2) silty claystones; coarse, conglomeratic sandstones are predominant near the gradational contact with Canebrake Conglomerate. The sandstones weather to an orange-brown color (5 YR 5/4 to 10 YR 7-8/1-3) which, from a distance, may resemble the colors of "C" sequences, but up close the distinction is obvious.

In much of the 011a member, parallel or subparallel lamination with low-angle truncations, low-angle cross-stratification, ripple lamination, and small-scale scour-and-fill structures (commonly gravel- filled) are typical. Debris-flow beds, consisting of poorly sorted sand, granules and pebbles in a dark clay-rich matrix, occur locally but are relatively rare. Bedding ranges from thin to very thick and is typically lenticular. "L"-suite sequences show local evidence of both upward-fining and upward-coarsening cyclicity, but in general exhibit much less consistent organization than do "C"-suite sequences of the

011a and Diablo members.

In the lower part of the 011a member, well exposed in Layer

Cake Wash (Figures 5.11, 5.12), lower Sandstone Canyon, and adjacent parts of Fish Creek Wash, "L"-suite sandstone beds contain abundant large-scale (0.5 to 2 m) tangential cross-stratification. In many cases 229 these sandstone sequences contain mud partings, mud-cracked strata, and angular, plate-shaped intraclasts derived from cracked mud layers. In Layer Cake Wash, a number of upward-coarsening and upward-thickening cycles were observed (Figures 5.11, 5.12). The abundance of large-scale cross-stratification generally increases upward in individual upward- thickening cycles.

Paleocurrent measurements were made from foreset orientations in tangential cross-stratification in "L" sandstone beds in this area

(fig. 5.15). These data show a broad dispersion, but a strong dominance of easterly directions, which contrasts with the south to southeasterly directions of "C"-suite palocurrents in the 011a member (Figure 5.1).

Paleontology

Observations about the Diablo member generally apply to

"C"-suite sequences of the 011a member. Silicified logs with discoloration haloes are common in "C"-suite sandstone bodies of the 011a member, as in the Diablo member. Carbonized palm fronds were found in the base of one upward-fining "C" sandstone body exposed in 011a

Wash. However, no reworked marine molluscs were found in the 011a member.

Woodard (1963) showed only one vertebrate locality in his

"upper submember" of the Diablo (011a member of this study), and did not discuss it specifically. However, camel tracks have been found along Arroyo Seco del Diablo in the vicinity of Camel Ridge (plate 4), in "L"-suite sandstones of the 011a member (P. Remeika, 1985, oral 230

20%

n=124

01la 'ILN Sands trough & tabular foresets

Figure 5.16. Paleocurrent rose for "L"-suite sandstones, 011a member, from Layer Cake Wash, Sandstone Canyon, and adjacent areas in Fish Creek Wash. All data are from medium to large-scale cross-stratification. 231 communication). These camel tracks are typically smaller (-10 cm long) than camel tracks found in the Jackson Fork member. Small (-1 cm) burrows, presumably of terrestrial invertebrates, are locally common in "L"-suite sandstone beds of the 011a member, for example in Sandstone Canyon. They are mostly vertical to diagonal, and are expressed by the disturbance of lamination. Similar burrows were

found in only one place in "C"-suite sandstone (also in Sandstone

Canyon), where disturbance of heavy-mineral laminations indicated their

presence. Thus, they may be common but cryptic in "C"-suite sandstones.

Lithologically equivalent units

A similar transitional facies between the Diablo and Canebrake

occurs on the eastern and southern flanks of the Santa Rosa Mountains

(Plate 3), as discussed previously. Dibblee (1954, 1984) included this

facies within his Palm Spring Formation, whereas Hoover (1965) recognized it as a separate unit and designated it "Facies B" of the

Canebrake Conglomerate (Figure 1.11). Hoover's map and stratigraphic descriptions provided the most detailed information available on this

011a equivalent in the Santa Rosa Mountains; his map has been incorporated in Plate 3.

Other occurrences of this alternating-provenance facies were observed by the author in regional reconnaissance, and were mapped as

011a member (Plate 3). In the western Borrego Badlands it grades upward into the Borrego Formation (Figure 5.2). In Borrego Mountain Canyon

(Borrego Mountain 7.5' Quadrangle), it occurs between the underlying

Canebrake Conglomerate and the overlying Diablo member. 232 In the Sierra de los Cucapas, the "Palm Spring Formation" described by Barnard (1966), which is essentially equivalent to the Diablo member of this study, locally rests on basement, and grades laterally and upward into the locally-derived "Canebrake Formation" over 10 to 30 m. Where the Palm Spring rests on basement, a basal facies is reportedly present, consisting of relatively immature sandstone and siltstone, and local lithic sandstone with local basement clasts (Barnard 1966). These basal and transitional facies are apparently analogous to the 011a member.

Canebrake Conglomerate (PQc)

The Canebrake crops out along the southeastern flank of the

Vallecito Mountains, and in the vicinity of Canyon Sin Nombre (Plates

3, 4). It was named by Dibblee (1954), who designated a type section at the southeastern base of the Vallecito Mountains, -5 km west of Fish

Creek Wash, and just west of the present study area. The Canebrake was described briefly by Woodard (1963), but was not examined in detail in the present study. It grades laterally into the 011a and Hueso members; the boundary is loosely defined by the change in dominant lithology from conglomerate to sandstone. The basal contact of Canebrake with plutonic basement was remapped along the southern flank of.the Valle- cito Mountains between Hapaha Flat and Jackson Fork. This demonstrates that the contact is depositional, and typically a high-angle nonconformity (butress unconformity). Woodard (1963) reported that "west of the type area the formation laps onto, or is faulted against bedrock" and the "conglomerate beds overlap unconformably onto granite bedrock south 233 of Sweeny [sic] Pass Road in the Jacumba Mountains." Woodard (1963) indicated a maximum thickness of -2100 m in the type area. In his

Canebrake Conglomerate he included units that interfinger with or grade laterally into members of the Imperial Formation; these include the

Jackson Fork and part of the Stone Wash members of the present study.

The Canebrake is typically a gray to grayish-tan, crudely bedded, polymictic, pebble to boulder conglomerate. The larger clasts are generally subrounded to subangular, and of predominantly granitoid

composition but with subordinate gneiss clasts and other minor con-

stituents such as amphibolite. The matrix is typically a poorly sorted,

angular, feldspathic coarse sandstone. Clast composition of the Cane-

brake, where examined along the southern flank of the Vallecito

Mountains east of Hapaha Flat, is distinctly different from the under-

lying basement, which is entirely plutonic. This suggests a larger

source area for the conglomerate than just the immediately adjacent mountains.

Compared to the typical Elephant Trees fanglomerate exposed in

Split Mountain Gorge, the Canebrake is distinguished by (1) a paucity

of dark-colored debris-flow beds and sharp bedding contacts; (2) poly- micitic clast composition; (3) general lack of reddish coloration

(although pale red beds were observed locally in the Canebrake

Conglomerate in Blackwood Basin); and (4) lack of evidence for syndepositional faulting. While the stratigraphy of the Canebrake is mostly quite monotonous, unusual bedding was observed at two locali- tites: (1) a -20 m thick bed of monomictic breccia in a tributary 234 canyon of upper Sandstone Canyon, composed of granitoid plutonic clasts; and (2) a thick (>20 m?) megabreccia bed with a dark green matrix in the canyon at the head of Blackwood Basin Wash (locality BW-m-11), similar to the boulder beds of Split Mountain.

Lithologically equivalent units

The Canebrake Conglomerate described from the southern Santa

Rosa Mountains (Dibblee 1954, 1984, Hoover 1965) is essentially the

same as that described from the FCV area. As in the FCV area, the

Canebrake grades laterally into an 011a-like facies, which grades in turn into a Diablo-like facies.

The Canebrake is also extensively exposed in Borrego Mountain,

in contact with crystalline basement (Dibblee 1984). Colors are highly

variable, but reddish-brown predominates, in contrast to the more

typical grays and olives of other Canebrake exposures. In a west-

dipping sequence exposed in Borrego Mountain Canyon (Borrego Mountain

7.5' Quadrangle), conglomerate grades upward into a sandy, "L"-suite

facies, which in turn grades upward into the 011a member, in turn

overlain by the Diablo member. Because of the regional west dip, the

Diablo member crops out west of the Canebrake, a reversal of the usual

outcrop pattern (Plate 3).

In the Sierra de los Cucapas, the "Canebrake Formation"

described by Barnard (1968) overlies basement or grades laterally and

downward into the "Palm Spring Formation" (Diablo equivalent). The

Canebrake reportedly consists of boulder conglomerates near basement

contacts, but grades laterally into conglomeratic arkose within 2 to 3 235 km. These conglomerates range in color from deep red (studied by Walker 1967) to reddish brown to drab brown (Barnard 1968). Schultejann (1984) described a unit of unconsolidated megabreccia overlying a low-angle detachment surface at Yaqui Ridge (Plate 3), and overlain unconformably by the Borrego (?) Formation. She suggested a correlation of this unit with the fanglomerates of Split

Mountain (Anza Formation of Woodard 1974). However, because this megabreccia is unconsolidated and does not underlie Imperial Formation, correlation with the Canebrake seems more reasonable.

Other basin-margin fanglomerates have been described from various localities around the Salton Trough; these units appear to be sedimentologically similar to the Canebrake. Only the stratigraphic relations are noticeably different:

1. Ocotillo Conglomerate (Dibblee 1954, 1984), a large tongue of conglomerate and sandstone in the Borrego Badlands and San Felipe

Hills, overlies the lacustrine Borrego Formation, underlies the lacustrine Brawley Formation, and pinches out toward the east (Figure 5.2, Plate 3).

2. The Painted Hill Formation near Whitewater Canyon overlies the Coachella fanglomerate and Imperial Formation (Proctor 1968).

3. Fanglomerates and conglomeratic sandstones in the Indio

Hills have been variously referred to as the Mud Hills Series (Free

1914), Indio Formation (Buwalda and Stanton 1930); Palm Spring Forma- tion (Dibblee 1954, Proctor 1968), and Thousand Palms and Indio Hills

Formations (Stotts 1965). 236 In summary, the working definition of the Canebrake as opposed to other fanglomerates is based primarily on stratigraphie relations. Thus, the essential characteristics of the Canebrake are onlap on crystalline basement and lateral interfingering with and gradation into nonmarine "L"-suite sandstone units such as the 011a and Hueso members.

Hueso member (PQh)

This unit consists of a thick (up to 1.5 km) sequence of

"L"-suite sandstones exposed north of Carrizo Wash in the Vallecito Creek, Arroyo Hueso, and Arroyo Tapiado drainages (Figure 5.1), and south of Carrizo Wash in the vicinity of Sweeney Pass and Canyon Sin

Nombre (Plate 4). The Hueso member, named by Woodard (1963), is known primarily for its rich vertebrate fauna (Vallecito Creek fauna of Downs and White 1968). In a considerable part of its outcrop area, this sequence is virtually undisturbed structurally; stratigraphic continuity has permitted magnetostratigraphic dating of the sequence

(Johnson et al. 1983). The outcrop area of the Hueso member includes the Palm Spring oasis which serves as the type locality for the Palm

Spring Formation. This is somewhat unfortunate as the Hueso member is a fairly local unit, whereas the Diablo member has a much more widespread distribution and is essentially synonymous with "Palm Spring Formation" as used in many areas. The Hueso member was not examined in detail in this study; Woodard's (1963) description is the most detailed available. Woodard (1963) also defined a "Vallecito member" which overlies the Hueso member. The author was unable to determine the basis for differentiating this unit from the Hueso member, except possibly a 237 coarser and more conglomeratic texture. In this study, Woodard's Hueso and Vallecito members are not differentiated, and are combined into the Hueso member (Figure 1.9).

The Hueso member in the FCV section overlies the 011a and Diablo members with sharp, conformable contact defined by the uppermost occurrence of "C"-suite sandstones, as discussed earlier. Along Arroyo

Tapiado and Arroyo Seco del Diablo, mudstones of the Tapido member intervene between the Hueso and Diablo members. The Tapiado grades laterally and upward into the Hueso but overlies the Diablo with sharp, conformable contact. To the northwest the Hueso and Vallecito members grade laterally into and interfinger with the Canebrake Conglomerate.

Woodard (1963) gave maximum thicknesses of -1000 m for the Hueso member and 105 m for the Vallecito member.

In Canyon Sin Nombre and vicinity the Hueso member represents a laterally gradational and interfingering facies between the 011a member and the Canebrake Conglomerate (Plate 4). It is distinguished from the

011a by the absence of "C"-suite sandstones and from the Canebrake by the predominance of sandstone over conglomerate.

Sedimentary facies

In upper Arroyo Tapiado, stratigraphically above the contact

With the 011a member, the Hueso member is virtually indistinguishable from "L" sandstone sequences of the 011a member. The color is predominantly olive to gray, and the dominant structures are subparallel lamination, low-angle cross-stratification, and small scour-and-fill structures. In fact, the Hueso could be described as the 011a member 238 without the "C"-suite sandstone sequences. The Hueso does contain some significant conglomeratic units, most notably a laterally continuous, ridge-forming unit described by Woodard (1963). The "Badlands Overlook" in the Arroyo Tapiado 7.5' Quadrangle is on this ridge. Higher in this sequence, in Arroyo Tapiado, lenticular, upward-fining cycles were observed. A typical cycle is <10 m thick and consists of a basal light buff, trough cross-stratified coarse sandstone, grading up through darker buff, fine sandstone, and into siltstone and mudstone with calcareous nodules.

Paleontology

Most of the vertebrate localities shown by Woodard (1963) for the FCV section are in the Hueso member. Downs and White (1968) listed the vertebrate fauna of the Hueso member (their Vallecito Creek fauna). They reported Megalonyx, cf. Nothrotherium, Sylvilagnus, Lepas,

Geomys garbanii, Stegomastodon, Equus, Smilodon, cf. Camelops,

Tanu lama, Tetrameryx, and Capromeryx as the dominant constituents, and assigned an Irvingtonian to latest Blancan age to this fauna.

Downs and Woodard (1961) and Woodard (1963) reported local occurrences of Ostrea vespertina and Anomia subcostata in the Hueso member. While these localities were not examined in the present study, it is suspected that these were reworked from the Imperial Formation

(Deguynos member), as in the Diablo member. 239 Lithologically equivalent units

The sandy facies of the Ocotillo Conglomerate ((k)) of Dibblee (1954, 1984) exposed in the Borrego Badlands in the vicinity of Fonts Point (Plate 3) is an "L"-suite unit that is lithologically similar to the Hueso member, and it occupies a somewhat analogous stratigraphic position (Figure 5.2). Age relationships between the ostensibly Pleis- tocene (Dibblee) Ocotillo and the Plio-Pleistocene Hueso member have not been established. Other units discussed under lithologic equivalents to the Canebrake Conglomerate have sandy "L"-suite facies that are also similar to the Hueso member.

Tapiado member (Pt)

The Tapiado member is essentially a lens of olive claystone with siltstone and sandstone, within the lower part of the Hueso member. Woodard (1963) reported the presence of freshwater gastropods and ostracodes, turtle fragments, and fish spines in this unit. Laterally continuous, parallel bedding is characteristic of the Tapiado (Figure

5.9, 5.10), in contrast with the more lenticular bedding geometries of the surrounding units. The lower contact with the Diablo and 011a members is sharp and marked by a striking color change. Sandstones and siltstones in the Tapiado are biotite-rich, and clearly of the "L" suite, in contrast with the dominant "C" suite of the underlying units. Woodard (1963) took as a lower contact the "fossil-bearing bed" at AT-m-16. However, as discussed before, this bed is actually a basal lag of gravel and reworked Imperial fossils in an upward-fining cycle in the Diablo member. The upper and lateral contacts of the Tapiado are 240 gradational with the Hueso member and were not explicitly defined by

Woodard (1963) nor will a definition be given here. For mapping, predominance of claystone and siltstone and thin, even bedding were used to distinguish the Tapiado member from the Hueso member, and the

absence of "C"-suite sediment distinguish it from the underlying Diablo

member. As mapped here, it is -200-250 in thick.

Lithology of the Tapiado grades from massive claystone to

interbedded sequences of claystone, siltstone, and very fine to fine

sandstone. Ripple lamination, particularly climbing-ripple lamination,

is the dominant sedimentary structure in the siltstone and sandstone

beds (Figure 5.9). A less common but very striking feature is the

presence of convolute lamination within some sandstone beds. This

convolution typically consists of a series of evenly spaced "rolls"

within a bed; the vertical scale is generally less than 0.5 m. This

style of convolute lamination is quite different from that of the

Diablo member, which occurs in massive sandstone bodies, is of larger

scale, and is more chaotic. In Diablo Canyon (DC-7b-227-230, Figure

5.9), large, evenly spaced, syn-depositional load casts were observed. Woodard (1963) mentioned the presence of limestone beds and

lenses within the Tapiado member. These limestones were not observed in

the present study. However, tuff beds were observed in a few places

along Arroyo Tapiado; these may actually constitute a single ash bed

exposed in different localities. Johnson et al. (1983) obtained a

fission-track age of 2.3 + 0.4 Ma on zircon from this tuff; this age

was used to calibrate their polarity-reversal sequence. 241 Lithologically equivalent units

The Tapiado member appears to be a unique, local occurrence.

Other lacustrine units occur north and east of the study area, and have been placed in the Borrego and Brawley Formations (e g. Dibblee 1984).

The Borrego and Brawley Formations, unlike the Tapiado, are dominated

by "C"-suite siltsones and sandstones, with very little mica, although

tongues of "L"-suite sandstone do occur in the Borrego Formation

(Figure 5.2). The Borrego Formation grades down very gradually into the

underlying Diablo member, whereas the Tapiado member overlies the

Diablo with a sharp contact defined by a striking color contrast.

Borrego (Pb) and Brawley (Qb) Formations

The Borrego does not occur in the FCV area, but it is important

in regional stratigraphy and was examined in this study in the San

Felipe Hills (Arroyo Salada and Tule Wash) and Borrego Badlands

(Hills-of-the-Moon Wash). The unit was named by Tarbet (1944) and was

described by Dibblee (1954) as light gray claystone interbedded with

sandstone and containing a "lacustrine fauna of minor molluscs, ostra-

cods, and rare Foraminifera." Dibblee (1984) later described the unit

as similar in appearance to the Imperial Formation but without the

"yellowish tan color and the oyster reefs".

In fact, there is little danger that the Borrego would be

mistaken for the Imperial Formation. While the massive gray claystones

are similar to those of the Imperial, the siltstone and sandstone beds

are pale orange and in some outcrops these are the dominant lithology.

As in the Tapiado member, ripple lamination is the dominant sedimentary 242 structure, but many sandstone and siltstone beds appear structureless, probably because of the textural uniformity and paucity of mica or heavy minerals to accentuate lamination. Also, like the Tapiado member, bedding is typically even and parallel. Interbedded Borrego sequences are distinct from claystone-siltsone-sandstone sequences of the Diablo primarily because of claystone color and bedding geometry.

The basal contact of the Borrego with the Palm Spring Formation is gradational; in fact, there is a thick transition zone (Figure 5.2) in which Borrego-like sequences alternate with Diablo-like sequences with reddish claystone and massive, upward-fining sandstone sequences.

This transitional unit is well exposed along Arroyo Salada and Tule

Wash, west of U.S. 78 (Kane Spring NW and Truckhaven 7.5' Quadrangles).

Mapping of the Diablo-Borrego contact is therefore arbitrary unless more explicit criteria can be developed.

The Borrego is succeeded by another lacustrine unit, the

Brawley Formation. These units are separated by a distal sandstone tongue of the Ocotillo Conglomerate (Dibblee 1984, Wagoner 1977); this tongue pinches out to the east and the two lacustrine units become indistinguishable. Dronyk (1977) and Wagoner (1977) examined sandstones of the upper Borrego and lower Brawley in the San Felipe Hills.

"C"-suite sandstones predominate, but paleocurrent directions based on small-scale ripple lamination are dominantly northeast to southeast.

Paleontology

Arnal (1961) briefly mentioned microfaunas of the Borrego and

Brawley Formations in his paper on the Salton Sea, but did not give 243 collecting localities. For the Borrego, he reported a mixed lacustrine and brackish-marine foraminiferal fauna dominated by Elphidium gunteri, with Nonion saltonensis, Streblus sobrinus, and S. tepidus, plus Chara oogonia. The partially marine fauna was explained in terms of gradual freshening of the Salton Trough basin following deposition of the

Imperial Formation. For the Brawley, he reported a strictly freshwater

lacustrine microfauna. The modern Salton Sea reportedly contains some

marine foraminifers, including miliolids, Nonion, Anomalina, Cibicides,

Rotalia, Elphidium, and Quinqueloculina, but he attributed these to

accidental introduction by man after the Colorado River was diverted. CHAPTER 6

INTERPRETATION OF FISH CREEK-VALLECITO SECTION

Overview

Most of the FCV and adjacent Neogene sections consists of two discrete bodies of sedimentary rock, distinguished by provenance and characterized by different paleocurrent directions. The "C"-suite sedimentary body is uniformly fine-grained, is clearly derived from the Colorado River (Chapter 2), and is dominated by southerly paleocurrent directions throughout the section (Figure 6.1). Sedimentologic observations supplemented by foraminiferal paleoecology (Quinn and Cronin

1984) demonstrate that this sedimentary body consists of an upward- shoaling and upward-freshening marine sequence grading into a homogeneous nonmarine unit (Figure 6.1). Henceforth, this sedimentary body will be referred to as the deltaic sequence (Figure 6.2). Colorado

River-derived turbidites of the Wind Caves member could arguably be included in the deltaic sequence, but they are discussed in this chapter with the pre-deltaic marine units. The term "deltaic" is used here in the sense of a body of sediment deposited subaerially and subaque- ously by a, river meeting a standing body of water. Thus, it carries an implication of common provenance and can encompass a wide variety of marine and nonmArine paleoenvironments.

244 245

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n•••4 • C:14 246 247 The "L"-suite sedimentary body is predominantly coarse-grained, consists primarily of arkosic sandstone and conglomerate, is in depositional contact with crystalline basement, and could clearly be derived from relatively local basement sources (Chapter 2). This sedimentary body is characterized by rapid lateral facies changes and easterly to northeasterly paleocurrents. It underlies, overlies, and interfingers with the "C"-suite deltaic sequence; therefore, the "L"-suite body can be loosely divided into pre-deltaic, syn-deltaic, and post-deltaic units. In this chapter syn-deltaic and post-deltaic "L"-suite units are discussed together as basin-margin alluvial units. Lacustrine units are discussed separately.

Further subdivisions of the pre-deltaic sequence are also useful. The association of nonmarine fanglomerates (Elephant Trees member) with syndepositional normal faulting (No Return fault) and alkalic to tholeiitic basalts (Alverson formation) is characteristc of extensional tectonic regimes (i. e. continental rifting), and these units are referred to as the svn-rift nonmarine sequence. Clast composition and paleocurrent directions of the underlying Red Rock member

(Kerr 1982, 1984) suggest extra-local derivation, at least in part, and therefore regionally integrated drainage. The Red Rock member is therefore regarded as largely pre-rift, although Red Rock deposition apparently overlapped with the onset of volcanism, as evidenced by basalt clasts in the Red Rock member (Kerr 1982, 1984). The end of continental rifting is more difficult to document than the onset, and extensional tectonics have probably been continuously active in the 248 Salton Trough since that onset; the designation "post-rift" is therefore not used.

The change from nonmarine to marine deposition permits further subdivision of the syn-rift sequence, and recognition of a pre-deltaic sequence. marine As mentioned before, the Wind Caves member is included

in this sequence. The Latrania member, which includes both marine and nonmarine facies, is discussed with the pre-deltaic marine units. The

Fish Creek Gypsum is somewhat problematic in this scheme, in that it

interfingers with apparently nonmarine units and contains no sedimento-

logic indicators of marine deposition, yet its stratigraphic position suggests that it was a direct precursor to marine inundation. There-

fore, it is discussed separately under pre-marine evaporites.

Fence diagram

Stratigraphic relationships in an area of rapid lateral facies changes are more readily visualized on a fence diagram (Figure 6.2) than on a geologic map. Construction of a three-dimensional diagram from essentially two-dimensional outcrop exposure requires a considerable amount of projection, and the result is necessarily somewhat artificial. As discussed in Chapters 4 and 5, most units of the deltaic sequence (especially Mud Hills, DeGuynos, Camels Head, and Diablo) are laterally persistent, in contrast to the "L"-suite units.

To provide a reference frame, the DeGuynos, Camels Head, and

Diablo members were projected laterally and depicted as isopachous units (this assumption is convenient but mostly unverifiable). Thick- ness variations in the Mud Hills member are fairly well established, 249 and are depicted on the diagram. Other stratigraphic units are shown in relation to the deltaic sequence as inferred from field relationships. Thicknesses of the pre-deltaic and post-deltaic units are depicted to

honor local thickness estimates as discussed in the stratigraphie descriptions.

Proximality and paleotopography

Lapout, proximality, and paleocurrent directions are important

for establishing the nature of basement paleotopography and deposition-

al paleoslopes. In the FCV area, the major basement high was clearly

toward the west, based on lapout of conglomerates on the southeastern

flank of the Vallecito Mountains; this basement paleo-high apparently

constituted the western flank of the sedimentary basin (Figure 6.2).

Proximality directions for the Latrania member in the Coyote Mountains

(Figure 4.2) are consistent with a major paleo-high to the south and

west. Lesser, local paleo-highs, such as Boundary Mountain (Figures

3.1, 6.2) and localities within the Coyote Mountains (S. M. Kidwell

1986, written communication), can also be recognized, but these were

apparently not major sediment sources. Basement paleotopography

probably included actively growing fault scarps, particularly in the

early stage of basin formation, but only one such syndepositional fault

(No Return fault) was identified in the FCV area. Average paleoslopes

of "L"-suite depositional surfaces were dominantly toward the east, as

inferred from paleocurrent directions, or away from the steep basement

topography of the western basin margin. These east-dipping paleoslopes

Mrsded into south-sloping depositional surfaces of "C"-suite sediments. 250 The coarser grain sizes of the "L" suite suggest that the east-dipping

paleoslopes were probably steeper than the south-dipping "C"-suite slopes.

Basement paleotopography was substantially different from the

present topography; in other words, modern topography does not repre-

sent exhumed paleotopography. This point has apparently not been fully

appreciated by some earlier investigators. For example, Stump (1972)

portrayed the paleogeography of early marine deposition in terms of

ancestral Coyote and Fish Creek Mountains. Similarly, Pappajohn (1980)

and Kerr (1982) discussed source areas for coarse clasts in terms of

present-day mountain ranges. Interpretation of coarse-clast provenance

is complicated by the possibility that some possible source areas are

now buried, such as the basement under the Vallecito basin, which was

probably exposed in the Miocene. In contrast to the present study area,

facies relationships in the Bouse Embayment (Metzger et al. 1973;

Metzger and Loeltz 1973) and Yuma area (Olmsted et al. 1973) suggest

that Neogene paleotopography in those areas was grossly similar to

Present-day topography.

Stratigraphy of "L"-suite units is characterized by lateral and

vertical hterogeneity and lack of a regionally persistent succession.

In the "L"-suite units, rapid lateral facies changes, particularly from sandstone to conglomerate, and lapout onto crystalline basement, suggest that alluvial fans were a major paleotopographic element along the western basin margin, similar to the present margins of the Salton

Trough and Gulf of California. Alluvial-fan facies in different 251 settings grade laterally into marine facies or "fan deltas" (Latrania,

Stone Wash, Lycium, Kamia Creek, and Jackson Fork members), fluvial facies (011a and Hueso members) and lacustrine facies (Tapiado member).

In contrast, "C"-suite units constitute a laterally persistent marine-to-nonmarine stratigraphic sequence that can be recognized not only in the FCV-CIA area, but also in the Coyote Mountains, Yuha Basin,

San Felipe Hills. A similar sequence also occurs in the subsurface of

the San Luis and Fortuna Basins (Chapters 4 and 5). Therefore, paleo-

surfaces of the ancestral Colorado delta were regionally extensive and

reached from the eastern to the western side of the Salton Trough.

Age control

The following constraints on ages of the FCV section and

adjacent sections were discussed in the stratigraphic descriptions and

are summarized here.

1. K-Ar dates on basalt flows of the Alverson Formation. These

range from -14 to -22 Ma, or early to middle Miocene. The basalt flows

predate marine transgression, represented by the Latrania member.

2. Biostratigraphy of marine macroinvertebrates in the Latrania recent member. Both Miocene and Pliocene ages have been suggested, with resolution is workers leaning toward Pliocene, but biostratigraphic

inadequate to unequivocally distinguish upper Miocene from lower Plio- macrofauna is cene. Part of the problem is that the Gulf of California age provincial, with strong affinities for Caribbean faunas, and

assignments have been based largely on comparisons with those for the Gulf, faunas. An independent molluscan biostratigraphy 252 calibrated with radiometric ages on volcanic strata, is now being developed, but has not yet been applied to the Salton Trough (Judith

Terry Smith 1986, oral communication).

3. Foraminiferal biostratigraphy of "C"-suite marine units in

the Fish Creek section. Stump (1972) reported planktonic zones N19,

N20, and N21 in the Imperial Formation. Based on his stratigraphie

descriptions, these zones apparently correspond to the Wind Caves, Mud

Hills, and Deguynos members of this study, respectively. The age of the

Lycium member is probably either late Miocene or early Pliocene, but

biostratigraphic resolution is inadequate due to the paucity of plank-

tonic foraminifers (Stump 1972).

4. Vertebrate biostratigraphy of the Palm Spring Formation in

the FCV section. Downs and White (1968) recognized Blancan and Irving-

tonian land-mammal ages to be represented, which would indicate a

Plio-Pleistocene age.

More precise age control on the FCV section was recently provi-

ded by magnetic-polarity chronology by Johnson et al. (1983) (Figure et 6.3). This study was based on samples from an earlier study (Opdyke

al. 1977) which were magnetically re-cleaned after a residual secondary

moment magnetization was discovered. The polarity chronology was an calibrated by a fission-track age of 2.3+0.4 Ma on zircons from

air-fall tuff in the Tapiado member exposed in Arroyo Tapiado. are: (1) 4.3 Important ages reported by Johnson et al. (1983) the base of the Ma for the base of the Imperial Formation (presumably al. used Woodard's 1974 Wind Caves member of this study, as Johnson et 253

VGP LATITUDE 0 0 -90 -45 0 0 +45 0 +90°

0--0

PQh

4—

Pt volcanic TTTTT ash

0- Pd •

1—

MPiu

0—

Figure 6.3. Magnetic-polarity stratigraphy of Imperial and Palm Spring Formations in FCV section, modified from Johnson et al. (1983). Graph shows site-mean latitutes of virtual geomagnetic Poles (VGP). "Open symbols represent data whose polarity is not in doubt, but which is not deemed statistically significant (k<10)" (Johnson et al. 1983). 254 nomenclature); (2) -4.0 Ma for the Imperial-Palm Spring contact (presumably the Camels Head-Diablo contact of this study); and (3) -0.9 Ma for the top of the Palm Spring Formation (i. e. top of the Hueso

member). The Diablo-Tapiado contact occurs between polarity reversals at 2.5 and 2.9 Ma, based on a sample-location map from P. Remeika (1985, written communication). The author estimated an age of 2.8 Ma for this contact by linear interpolation of stratigraphic thickness;

N. M. Johnson (1985, written communication) concurred with this age estimate. Of these ages, the base of the Imperial is probably the least reliable, because of the relatively poor quality of data from the Imperial (Figure 6.3). Failure to recognize an interval of reversed

polarity would result in underestimation of the age for the base of the

sequence. For this reason, the age of 4.3 Ma is regarded as a minimum.

The Lycium member, the lowest marine unit in the FCV section, is older; its maximum age is still poorly constrained.

The age range of 4.0 to 4.3 Ma for the "Imperial Formation"

contradicts Stump's (1972) biostratigraphy. Only Zone N19, which Stump

recognized in the Wind Caves member of this study, falls within this

age range, while N20 and N21 are too young. This discrepancy is not

easy to account for, because Stump only mentioned the occurrence of

certain "index" foraminifers, but did not show the full stratigraphic

ranges of foraminifers in the Split Mountain section. Conceivably, some

of the problem may have to do with paleoenvironmental changes caused by progradation of the ancestral Colorado delta (Figure 6.1). 255 Magnetostratigraphy of the FCV section permitted Johnson et al. (1983) to calculate sedimentation and tectonic subsidence rates (Figure 6.4). They reported declining sedimentation rates from 5.5 to

0.5 mm/a, which approximately fit an exponential decay curve, except

for the lowest part of the section, where calculated rates are anomalously high. Corresponding tectonic subsidence rates (obtained by

dividing by isostatic amplification factor for sediment load) are 1.5 to 0.1 mm/a. An uplift rate of 5.9 mm/a within the last 0.9 Ma was

calculated, based on a stratigraphic thickness of 5 km (Johnson et al. 1983). If the additional stratigraphic thickness exposed at Split

Mountain is included, an uplift rate closer to 7 mm/a is obtained.

Pre-rift phase

Conglomeratic sandstones of the Red Rock member in part pre-date local volcanism and fanglomerate deposition, based on stratigraphic position (Figure 6.5), and are in part contemporaneous with volcanism, based on volcanic clasts within the Red Rock. However, Red

Rock lithologies are generally not interbedded with the Elephant Trees fanglomerates and Alverson volcanics, suggesting that Red Rock deposition mainly represents a distinct pre-rift phase. Lenticular bedding geometry, upward-fining cycles with basal conglomerates, and lack of interbedded fine-grained deposits are indicative of deposition in a bedload-dominated fluvial system. Kerr (1982, 1984), following the classification of Miall (1978), has interpreted these rocks in terms of braided-stream deposits of the Donjek and Platte types. Kerr (1982) 256 6.4

TIME (Ma B.P.)

5 FCV 4 section

3 _

2 exponential 1 decay

0 3' 2 TIME (Ma B.P.)

Figure 6.4. Cumulative sediment accumulation and sedimentation rates for the FCV section, based on magnetostratigraphy (Figure 6.4) Modified from Johnson et al. (1983). Sedimentation rates after -4 Ma approximate an exponential decay curve. 257 also interpreted large-scale foresets exposed in Red Rock Canyon as a product of eolian deposition.

More significant from a regional standpoint is the presence of extra-local metavolcanic clasts in Red Rock conglomerates (Kerr 1982), in contrast to post-Red Rock conglomerates which could all be locally derived. This suggests that the Red Rock member represents part of an

integrated drainage network. Confinement of the Red Rock to discrete

lenses between basement and syn-rift units (Figure 6.5) indicates that aggradational fluvial deposition was confined to valleys incised in

bedrock. Paleocurrent data suggest that the predominant paleoflow was northerly to westerly, although these data may primarily reflect the

local orientation of paleovalleys. Sedimentologic differences between the Split Mountain and CIA occurrences (Kerr 1982, 1984) could indicate that the two localities represent two discrete river systems in a drainage network, or that the Red Rock locality represents a smaller tributary and the Split Mountain locality a trunk stream. The cited directions are in reference to present geographic north, and could be affected by tectonic rotations (Chapter 8).

The presence of basalt clasts within the Red Rock member (Kerr

1982) indicates that volcanism may have begun in the nascent Salton

Trough earlier than in the immediate study area. Also, the Red Rock-

Alverson contact may be significantly diachronous, so that some of the

Red Rock volcanic clasts may be derived from laterally adjacent

Alverson flows. More thorough documentation of the stratigraphic

Position of volcanic clasts within the Alverson and radiometric dating 258 259 of such clasts might help to clarify details of the onset of rift-

related volcanism.

Syn-rift nonmarine phase

The syn-rift nonmarine phase (Figure 6.6) comprises the

Alverson volcanics, Elephant Trees fanglomerates, the Upper Boulder

Bed, and the Fish Creek gypsum.

Volcanism

Alkali olivine basalt and tholeiitic basalt, which constitute

most of the Alverson Formation (Gjerde 1982), are generally regarded as

characteristic of regimes of crustal extension (e. g. Martin and

Piwinskii 1972). Early to middle Miocene ages (-14-22 Ma) on the vol-

canic flows are therefore interpreted to date the onset of continental,

rifting in the nascent Salton Trough.

Brecciated flow tops, typical of aa flows, and general absence

of pillows (Ruisaard 1979), indicate that Alverson basalt flows were

mostly subaerial. Stratigraphic and geochemical characteristics of the

Alverson volcanics are quite localized. Although Ruisaard (1979)

attempted a subregional correlation, Gjerde (1982) demonstrated geo- and graphic localization of alkalic (Alverson Canyon and vicinity) with overlapping tholeiitic (CIA and Painted Gorge) geochemical suites

age ranges. This suggests the existence of multiple vents and magma

ambers, and limited lateral extent (a few km or less) of individual

flows. Geographic isolation of basalt flows from presumably correlative the fanglomerates, the paucity of non-volcanic clastic sediment within 260 261 Alverson Formation, and the lack of coherent internal stratigraphy in the Alverson all indicate that regional drainage was disrupted during the rift phase.

Fanglomerates

Kerr (1982, 1984) interpreted the lower Elephant Trees

conglomerate unit at Split Mountain Gorge in terms of an eastward-

prograding alluvial fan. Debris-flow and sheetflood beds dominate the lower part of the sequence, and tend to thicken up-section. In the

upper part of the sequence, channeled, clast-supported conglomerate

beds predominate over debris-flow beds. The overall facies sequence was

interpreted as distal fan, mid-fan, and channeled upper-fan facies.

Kerr (1982, 1984) regarded the Vallecito fault as the main syndepositional fault. However, the Elephant Trees fanglomerate does not

reach as far west as the Vallecito Fault; the No Return fault appears to be a more likely syn-depositional fault. The north-northwest trend of this fault is approximately aligned with the trend of volcanic necks reported by Ruisaard (1979) from the Alverson Formation in the CIA.

Source areas for the various fanglomerate units are somewhat problematic. Kerr (1982, 1984) has analyzed possible source areas in terms of present basement exposures. Two alternative possibilities should be considered, however: (1) basement rocks now covered by Plio-

Pleistocene strata may have been a source area during the Miocene, and (2) post-Miocene erosion of areas like the Vallecito Mountains with plutonic rocks may have stripped off metamorphic roof pendants. 262

From a regional standpoint, the Red Rock-Elephant Trees contact is important in that it marks a major drainage reorganization from north/northwest (Figure 6.5) to east/northeast (Figure 6.6). A drainage reversal was observed in the Puertecitos area of Baja California Norte, which occurred at approximately the same time (14-9 Ma) as a major episode of volcanism (Dokka and Merriam 1982). If, as seems likely, the

Elephant Trees and Alverson are coeval, the drainage reversal in the

FCV section may be directly analogous to that at Puertecitos.

Lower Boulder Bed

Robinson and Threet (1974), Kerr et al. (1979), and Kerr (1982,

1984) interpreted the LBB megabreccia as a subaerial landslide deposit,

and suggested transport on a cushion of trapped air. Clast composition

and the imbrication of clay seams support a source from the north or

northwest, probably from the area of the Vallecito Mountains (Kerr

1982, 1984).

The LBB has the following features in common with a general

model for subaqueous debrites (debris-flow beds) illustrated by Stow

(1986): (1) inverse grading of coarse clasts, with boulders concen- the base trated in the upper part of the bed; (2) intense shearing near described by of the bed; and (3) local thrust faults within the bed, as

Kerr (1982, 1984). According to Stow, such debrites include not only mud". Sharp muddy beds, but also "bouldery mass[es] containing little in debrites tops with large protruding clasts are reportedly common the LBB (Stow 1986), as observed in the LBB. These similarities between emplacement and subaqueous debris flows raises the possibility that 263 could have been under water, at least in part. Although the LBB clearly predates marine transgression, the possibility of local, ephemeral lakes cannot be discounted. The LBB was apparently deposited in a paleotopographic low, as suggested by the distribution of directly overlying Fish Creek Gypsum.

Unconformities

Local angular discordances occur within the lower nonmarine section, such as those described by Ruisaard (1979) above and below the

Alverson volcanics, and the one near the top of the Red Rock member in

Split Mountain Gorge. However, none of these appear to be very extensive; none could be traced throughout the study area. On trailing-edge continental margins, pronounced angular unconformities marking the change from active rifting to passive subsidence are commonly observed on seismic-reflection profiles (Bally 1981, Scrutton 1982); Falvey

(1974) referred to these as "breakup unconformitites". Because the western Salton Trough has behaved somewhat like a trailing continental margin, a similar breakup unconformity might be expected. Some possible explanations for the absence of a throughgoing unconformity are: (1) subsidence was rapid and sediment was continuously supplied to the basin, so there was no opportunity for complete erosional beveling; (2) block rotation associated with rifting was minor and local; or (3) marine inundation began while rifting was still underway, so post-rift subserial exposure was not widespread. Riba (1976) described intraformational unconformities exposed in the Spanish Pyrenees, in which

Ste-ply angular unconformities grade laterally over short distances 264 into lower-angle "progressive unconformities", which in turn grade into conformable surfaces. He referred to this association as "syntectonic unconformities"; this term may also apply to unconformities in the FCV area, although the tectonic setting is quite different.

Pre-marine evaporites

In the field, the Fish Creek Gypsum itself provides few clues

about the nature of its deposition, other than the exceptional purity of the deposit. However, stratigraphic context places some constraints on depositional models. Some of the features to be explained are: (1) the paucity of interbedded terrigenous sediment except near the base;

(2) the stratigraphic position, overlying entirely nonmarine deposits, and underlying entirely marine deposits; (3) the occurrence of gypsum as a single, large, discrete lens, rather than as multiple bodies or beds dispersed within marine or nonmarine strata; (4) the succession from gypsum to turbidites, without intervening brackish-water or shallow-marine deposits; and (5) the apparent absence of a typical sabkha facies tract or associated sedimentary structures such as replacement textures.

Similar pre-marine evaporites on a much larger scale are widespread in Mesozoic rift or extensional basins of the circum-Atlantic region (Evans 1978, Hardie 1984), and apparently represent a singular event in the evolution of such basins. Kendall (1984) has outlined a

"shallow-water, deep-basin" model for evaporite deposition that can explain most of the features of the Fish Creek Gypsum. Sea water is introduced into a topographic depression below sea level by seepage 265 through a sill, or by periodic overflow. Deposition can be very rapid, so that the necessary conditions have to be maintained only a short

period of time, and contamination by terrigenous sediment can be relatively minor. The thickness of the deposit is limited only by the depth

of the pre-existing depression, and does not depend on syn-depositional

subsidence. Once the sill is breached, the basin could be filled vir-

tually instantaneously. This could explain the fact that the Fish Creek

Gypsum is directly overlain by turbidites rather than by the typical shallow-water facies of the Latrania member.

Pre-deltaic marine phase

Pre-deltaic marine paleoenvironments in the study area (Figure 6.7) can be subdivided into (1) fossiliferous, locally-derived, shallow-marine deposits, (2) locally-derived subaqueous gravity-flow deposits, including sandstones and conglomerates, (3) the UBB, a subaqueous debris-flow megabreccia; and (4) Colorado River-derived turbidite sandstones. These facies do not correspond precisely to mappable stratigraphic units. Although some of these facies may grade into nonmarine deposits, indicators of paleoshorelines are difficult to find. Furthermore, although both "shallow" and "deep" facies are present, bedding geometries representing topset-foreset or shelf-slope stratification were not observed.

Fossiliferous shallow-marine facies

Fossiliferous facies of the basal Imperial Formation have been consistently interpreted by previous workers as shallow-marine, 266 267 littoral to neritic deposits on the basis of marine macroinvertebrates. Sandstone-to-conglomerate facies changes, as observed in Fossil Canyon, and lapout against basement, as observed in Andrade Canyon, suggest that small alluvial fans with marine fringes (i. e. fan-deltas) were the dominant depositional system of the Latrania member. Plutonic, metamorphic, and volcanic basement was locally exposed on the seafloor as well, as evidenced by lithophagid borings in marble and oysters

(Ostrea heermanni) resting directly on Alverson basalt (S. M. Kidwell

1986, written communication). Sedimentary structures in the Latrania

are conspicuous in their absence; this is largely attributed to

intensive bioturbation of sandstones and of sandstone matrix in conglomerates.

Dominance of biogenic over physical structures is usually characteristic of low-energy environments, but this presents a paradox, because sufficient energy was available to transport pebbles and cobbles as well as large bioclasts such as coral heads and thick-walled oyster shells. The most likely explanation is that coarse clasts were mobilized and reworked during infrequent high-energy events. Such events could include both storm-induced flash floods on alluvial fans, which could have prograded sand and gravel over fossiliferous beds, and storm-induced marine waves, which could have reworked and mixed terrigenous clasts and bioclasts. Little direct evidence was seen for tidal influence, but a high tide range could account for a wide zone of interaction between terrestrial and marine processes. 268 "Turbiditoid" facies

Locally-derived, arkosic sandstone and conglomerate, with sedimentary structures indicative of deposition by subaqueous, sediment gravity-flows, comprise the Kamia Creek member, Lycium member, distal Stone Wash member, and the lowermost part of the Wind Caves member (Plate 5). Facies relationships with shallow-marine to nonmarine conglomeratic units (proximal Stone Wash and upper Elephant Trees members), east-directed paleocurrents, and nearby westward lapout

(Figure 6.7) suggest that the turbidites are a distal facies of small, high-gradient alluvial fans which built out into fairly deep water from the western basin margin. The conventional model of a submarine fan fed by a single submarine canyon is probably not applicable here; the term

"slope apron" of Stow et al. (1985) may be more appropriate. Unfortun- ately, the distribution of outcrops is inadequate to show the complete proximal-to-distal (west-to-east) facies tract, so facies patterns must be inferred indirectly. Structures in the Lycium, Kamia Creek, and distal Stone Wash members range from classical turbidites to conglomerates with few process-diagnostic attributes, but stratigraphie context and facies relationships suggest deposition by subaqueous sediment-gravity flows. The general term "turbiditoids" is suggested for this facies association.

Mechanics and classification of coarse-grained, high-density, subaqueous gravity-flows, or "fluxoturbidites", have been the subject of a large body of literature (e. g. Middleton and Hampton 1976, Carter

1975, Lowe 1982, Pickering et al. 1986). However, relatively few facies 269 models have been developed specifically for turbidites and gravity flows associated with alluvial fans in rift basins (Surlyk 1984, Stow

1985), and well-documented analogs for marine deposits of the Split

Mountain area system are apparently rare. Bedding types were classified according to the scheme of Mutti and Ricci-Lucchi (1978, Howell and

Normark 1982), but such classification should be subordinate to

detailed section logs and stratigraphic context in making paleoenvironmental interpretation. Interpretation of Mutti and Ricci-Lucchi facies

in terms of their generalized submarine-fan model is probably not

appropriate here. A field classification of conglomerates in subaqueous

gravity-flow deposits proposed by Pickering et al. (1986) is concep-

tually useful, but is difficult to apply in the present study, in part

because of the difficulty of locating boundaries between flow units.

However, a model presented by Clifton (1984) for proximal-to-distal

changes in mixed sand-gravel density flows (Figure 6.8) may be

particularly applicable to this study.

Water depths and paleos lope gradients associated with these described depositional systems are not well constrained. Ingle (1974) "shallow the marine "Split Mountain" (Lycium of this study) fauna as this study) water", and interpreted the "lower Imperial" (Wind Caves of and Cronin microfauna as "upper bathyal" (>100 m) to "shelf". Quinn

(1984) interpreted their "middle Imperial" (also equivalent to Wind

Caves of this study) microfauna as "outer shelf". Such paleobathy-

metric designations are somewhat misleading, as true shelf-slope

morphology was probably not developed at this time, and the sedimentary 270

( . 4.) ..-1 cn 11) 0 CU •-• C. • H 0400 O c... cd O -•-I cs. O 1"4 Cf ri CD t) CO 4-) -H i.. G a.) d) 0 g E .0 0o 0 ✓...1 r-.1 .L-( to O• C g 0 i 0 cf) 0) 0. O "0 O - g .1.) c aS -0 :.> . G c L1) RI 0 c CL. ce a) 1. cd "C1 'cLo O 0 4-, •L-1 0 cn >-. a) 0.) 0 co O C) 0 0 I 0 c..) C.) 0 •.-1 ci-, ,-. S. o cd 0 S. 0 Cl.) Q) .9 O u) 4-) r-. X 4-. g "i' O O - O C.) 0 O cd 10 ... 4., 0 4.) cl) g cd S. -H +) 0 41 0 0.) 0 ›', V) • r ri Q) bo 5-1 CI = 04000 0) ' ...1 ;••1 n $•• .1.)O 4.)w O w •fft SW 0 4) ''"4 4' 01 ..0 0 0 .4.) Z .0 14 • c.) 0 --- cn 0..bd• 01 010 • 0 0") cc 0 L-4 . L. 271 facies at Split Mountain are completely uncharacteristic of shelf environments. However, the concensus on paleoecology seems to favor fairly shallow depths by turbidite standards, possibly even <200 m. Stratigraphic relationships indicate that the basin was quite small. The thick "turbiditoid" marine sequences at Split Mountain and Kamia Creek are bracketed by nonmarine units and the Mud Hills member. In adjacent areas, the comparable stratigraphic position is occupied by thinner sequences of fossiliferous, shallow-marine facies

of Latrania and proximal Stone Wash (along Coral Wash and east of Hill 1134, 4 km west of Split Mountain Gorge; east of Red Rock Canyon, 8 km

east of Split Mountain Gorge and 1 km west of Kamia Creek; and at

Barrett Canyon, 1 km east of Kamia Creek). Therefore, the maximum

northwest-southeast width of the deeper basin or embayment was only -12

km for Split Mountain and -2 km for Kamia Creek. Thickness variations

in the Mud Hills member may be an indicator of relative water depth of underlying units; the Mud Hills is thickest along Fish Creek Wash and

North Fork, above the Wind Caves member, and thinnest where it overlies shallow-marine facies of the Latrania and proximal Stone Wash members.

In the Lycium member, the association of Scolicia and other abundant horizontal burrows with Ophiomorpha and Thh...assinoides is atypical of turbidite ichnofaunas, and suggests a hybrid or intermediate between the Nereites and Cruziana ichnofacies of Frey and

Pemberton (1984). Occurrences of Ophiomorpha and Thalassinoides in turbidites, subaqueous gravity-flow deposits, and "deep-sea fans" have previously been reported from the Eocene of Spain (Crimes 1977) and the 272 Upper Cretaceous of California (Kern and War-me 1974, Bottjer 1981, Buck and Bottjer 1985).

Two such systems with different facies associations can be identified in the study area. In one facies association (units Meu, MPlt, MP11), typical nonmarine fanglomerates grade laterally into

debris-flow sandstone beds, and in turn into thin-bedded, laterally persistent turbidites. This facies association occurs just above the

stratigraphic level of the Fish Creek gypsum in Split Mountain, and

probably represents a relatively shallow, low-relief system. The second

facies association (units MPs, MPlu) constitutes a much larger volume,

and includes unusual features such as gravity-flow conglomerates and

large-scale, retrogradational sigmoid stratification (Figure 6.9).

The first facies association (units Meu, MPlt, MP11) consists of the following:

1. Fanglamerates. This is a typical association of alternating sheetflood and debris-flow conglomerate beds of the upper Elephant

Trees member. Clast size and bedding thickness generally decrease from northwest to southeast.

2. Transitional sandstone facies. This facies, which constitutes the transitional unit of the Lycium member, is exposed in Split

Mountain Gorge, and is thin to medium bedded, the thicker beds consisting of poorly sorted debris-flow sandstone beds. Thin lenses of gypsum occur locally in this facies. This facies lacks clear marine indicators. 273

274 3. Thin-bedded turbidites. The Fish Creek Gypsum is overlain by thin-bedded, fine to medium grained, laterally persistent (on outcrop scale) turbidites referred to previously as the "lower turbidite unit of the Lycium member. No evidence was found for a more distal facies, and these turbidites apparently occupied the same local topo-graphic depression in which the gypsum was deposited. The lack of intervening

fossiliferous marine beds suggests that a sub-sea level depression was rapidly flooded. In the Mutti and Ricci-Lucchi (1978) classification, these turbidites belong to Facies C to D. Bioturbation of bed tops is

typical for the Nereites ichnofacies (Frey and Pemberton 1984), although only one type of discrete trace fossil, Scolicia, was identi-

fied. Extensive bioturbation suggests that full marine, well-oxygenated bottom condiditons prevailed by this time. Sparse sole marks suggest a southerly transport direction.

The tops of facies 1 and 2 are marked by local evidence for wave reworking, including parallel lamination (SM-2-21, Plate 5) and tightly packed cobbles (both are exposed in Lycium Canyon). These may constitute a thin transgressive facies.

The second facies association (units MPs, MPlu) is volumi- trically dominant over the first and is much more extensively exposed; representative facies (except for proximal conglomerates) are illustrated in Plate 5. Extraction of a facies tract from these data is highly interpretive, as a complete tract of demonstrably time-equivalent facies is not exposed. 275 1. Proximal conglomerates and sandstones. This facies is the fossiliferous, proximal Stone Wash member, and is very similar to the

Latrania member. Characteristics were described under "fossiliferous

shallow-marine facies". Clast composition along Coral Wash is quite different from that of the distal Stone Wash member at Lycium Canyon,

so this is probably not a direct upslope facies equivalent. However, it may be representative of contemporaneous shallow-water facies.

2. Distal conglomerates. This facies is the distal Stone Wash member. The Stone Wash conglomerates exposed in Lycium Canyon tend to be crudely bedded and less well sorted than conglomerate beds of the underlying Lycium member. Some discontinuous partings of mottled silty sandstone separate succesive flow units; these may represent bioturbated flow tops although discrete burrows are generally not apparent

Interbedding of poorly organized conglomerates with turbidite sandstones is observed in Oyster Shell Canyon #3 (Figure 4.9), which supports a subaqueous interpretation. These conglomerates may have been deposited by dense sediment gravity flows, but the precise mechanism is not clear.

In the Lycium Canyon section (Figure 4.7, Plate 5) the gradation from Lycium member into Stone Wash member is best illustrated.

Conglomerates are most abundant above the UBB, and gravels are moderately to poorly sorted. Individual flow units are difficult to recognize; fine-grained parting between conglomeratic beds are generally discontinuous. As a result, grading in the coarse clasts would be difficult to detect. Farther down-section, and presumably more 276 distally, gravel clasts become less abundant but better sorted and on the average coarser grained, resulting in a strikingly bimodal mixture of large cobbles to boulders and sand. In the Lycium member, gravel clasts are sparse and scattered, and mainly limited to the uppermost part of the section. This progression of sand-gravel bedding types from proximal to distal is similar to the model developed by Clifton (1984)

for the Paleocene Point Lobos submarine-canyon fill in California (Figure 6.8). The main difference is that Clifton observed well-developed inverse grading of coarse clasts in the Point Lobos; such inverse grading is only locally and poorly developed in the Stone Wash member.

3. Channeled slope facies. This facies is relatively minor in outrop extent, and is only locally developed. It is best seen in Split

Mountan Gorge below the UBB, where it contains three bedding types. The first is thinly bedded, heavily bioturbated medium to fine grained sandstones with abundant ripple lamination; it is classified here as

Mutti and Ricci-Lucchi's (1978) facies E. The second bedding type consists of medium bedded, coarser grained graded sandstones with cobbles and boulders, occurring singly, dispersed within beds, and

Clustered (boulder nests). The third bedding type consists of small, meter-scale channels filled with graded sandstone beds. This facies may represent local development of relatively fine-grained slopes cut by small channels through which coarser sediment was bypassed. Larger- scale channeling, such as would be expected for upper fan environments in the Mutti and Ricci-Lucchi model, was not observed, however. 277 4. Sandy turbidite (slope apron) facies. This facies is extensively exposed and makes up the bulk of the Lycium member. Vertical sequences are dominated by medium-bedded coarse-grained turbidites and (Mutti Ricci Lucchi facies B), in some cases with well-developed Bouma sequences (facies C). Thinner-bedded sandstones with ripple lamination and intensive bioturbation (Facies E) are interbedded with

these. Coarser clasts are uncommon. The association of Facies B and E

is most characteristic of channel-levee complexes in upper (proximal)

submarine fan settings (Mutti and Ricci-Lucchi 1978; Howell and Normark

1984). The lower turbidite unit of the Lycium member may represent a

more distal facies of the same tract, but since it is not laterally

equivalent to the upper turbidite unit, this is not clear.

Much of the lenticular bedding geometry in the Lycium member

can be attributed to large-scale retrogradational sigmoid stratifi-

cation described in Chapter 4, and not to typical submarine canyon

geometries. The most diagnostic feature of this stratification style is

the truncation surfaces, which have a consistent sense of truncation

("down-cutting" in the down-current direction) (Figure 6.9). The sigmoid stratification has not, to the author's knowledge, been recognized in other turbidite sequences. This stratification style probably represents up-current migration of some large-scale, antidunelike bedform (Figure 6.9).

The Kamia Creek sequence (Plate 5) is similar to the facies tract just described. The uppermost facies consists of ungraded conglomeratic sandstone beds with a shallow-marine macrofauna. Below this 278 a sequence of is relatively thin and fine-grained graded beds. Below

this is the bulk of the section, characterized by coarse-grained graded beds similar to the upper turbidite unit of the Lycium member. In the Bouma classification, these are mainly AE and ABE beds. This facies sequence suggests a small-scale, eastward-prograding fan-delta system with a steep submarine slope.

Upper Boulder Bed

The Upper Boulder Bed marks a major interruption in the

sequence at Split Mountain and, judging by the extensive disruption and

truncation of the underlying Lycium member, must have considerably

altered the sea-floor topography. This change in the basin floor may have permitted the initial influx of "C"-suite sediments, which first appear a short distance above the UBB.

The Upper Boulder Bed megabreccia was clearly emplaced under water, probably as a debris flow; the upward gradation into sand at the head of Split Mountain Gorge suggests a late-stage change into a turbidity flow after the bulk of the sediment was emplaced. Elsewhere, the top of the flow is very irregular, with protruding boulders, as is mouton for submarine debris-flow beds (Stow 1986). Large intraclasts of the UBB have been reworked into overlying sandstone and conglomerate beds, possibly due in part to instability of the irregular top of the bed.

The source of the UBB is problematic. If fold axes are transverse to the paleoflow direction, they would suggest a southeasterly or northwesterly transport direction, and the sense of fold vergence and 279 undercutting exposed in Split Mountain Gorge suggest a southeasterly direction. However, the sense of overturning in a large number of folds does not show a significantly preferred orientation (Figure 4.11).

Alternatively, if fold axes on the average are parallel to paleoflow

directions, these directions could be northeasterly (similar to most "1"-suite units) or southwesterly.

Furthermore, the high content of metamorphic boulders is

incompatible with the plutonic lithologies exposed in the Vallecito

Mountains toward the northwest. For this reason, Pappajohn (1980)

preferred a source to the east, such as the northern Fish Creek Moun-

tains. Other possible explanations for the clast composition consistent

with a more westerly source are (1) the source area was to the south of

the Vallecito Mountains and is now buried by Pliocene strata of the FCV

section; (2) the source included metamorphic roof pendants in the

Vallecito Mountains which were subsequently removed by erosion; or (3)

the source area has been displaced by strike-slip faulting. This last

Possibility is the least likely, as major intervening strike-slip

faults have not been identified.

Wind Caves submarine fan.

"C"-suite turbidites of the Wind Caves members present a

striking contrast to the Lycium, Stone Wash, and Kamia Creek members, not only in sandstone provenance but in paleocurrents, bedding geome-

try, and sedimentary structures as well. Turbidite sandstones of the

Wind Caves member have a more distant source (Colorado River) than the underlying turbiditoid deposits, and the radial pattern of 280 paleocurrents (Figures 6.1, 6.7) suggests a true fan morphology. As the lower with turbidites, however, the complete proximal-to-distal (north-to-south in this case) facies tract is not exposed, and so there is no direct evidence for a feeder submarine canyon. Stratigraphic constraints (Figure 6.7) limit the width of this fan at Split Mountain to -6 km, and the downcurrent extent to <15 km. Thus, the Wind Caves fan is comparable in scale to the modern Navy Fan (Normark and Piper 1985), one of the smallest of well-studied modern submarine fans.

Vertical sequences of the Wind Caves member are upward-thinning

overall, which might suggest a retrogradational sequence. "C"-suite

sandstone beds in the lower Wind Caves member are very thick (up to 8 m), massive, graded, and lenticular, particularly near the confluence

of Fish Creek and North Fork. This facies is most similar to Facies B

of Mutti and Ricci Lucchi (1978) and probably represents a submarine

channel complex. Farther up in the section, bedding becomes thinner and more laterally persistent, with Bouma ABE, ACE, and ABCE beds. This upper sequence of the Wind Caves member is similar to Facies C (classical turbidites) of Mutti and Ricci Lucchi (1978), and is most characteristic of middle to outer fan settings.

One explanation for the upward-thinning character of the Wind

Caves sequence would be retrogradation of a submarine fan as subsidence exceeded sedimentation. However, in view of the overall progradational and upward-shoaling character of the "C"-suite sequence, this interpretation is suspect. An alternative interpretation is that the morphology of the basin floor changed during Wind Caves deposition, from 281

(1) a steep, high relief depression with a relatively deep axis occupied by a channel-levee complex, to (2) a shallower, lower-relief, non--channelized submarine fan or basin plain. This change in submarine topography could have resulted from a rapid influx of fine-grained

Colorado-derived sand. In other words, the vertical sequence of the

Wind Caves is attributed to a qualitative change of facies tract due to

shoaling and reduction of relief in the basin, rather than to simple

retrogradation of one facies tract caused by deepening of the basin.

Summary of pre-deltaic marine phase

Locally-derived marine sandstones and conglomerates represent a

variety of alluvial-fan/fan-delta environments. These alluvial fans

lapped out mainly toward the west, against high-relief basement topo-

graphy. Shallow-marine macroinvertebrates occupied the low-energy

marine fringes of alluvial fans, with sandy and gravelly substrates, as

well as submarine outcrops of volcanic and crystalline basement rocks.

Alluvial fans building into deeper water resulted in local accumula-

tions, at Split Mountain and Kamia Creek, dominated by sediment

gravity-flow deposits, mainly by high-density turbidity flows. Because have of the small size of these local basins, submarine gradients could

been quite steep without requiring great water depths.

Turbidites composed of Colorado-derived sediment were appar- basin. This ently deposited as a small submarine fan in a semi-enclosed

depositional system evolved from a channel-levee complex which presum-

ably occupied the axis of the bathymetric basin, to a low-gradient, 282 non-channelized fan as the basin was filled by the rapid influx of fine-grained sediment.

Deltaic sequence Stratigraphic sequences of deltas differ greatly in detail (e. g. Coleman and Wright 1975, Coleman and Prior 1982, Elliott 1986),

but a generalized profile can be described for fine-grained, prograda-

tional, marine deltaic sequences. The following generalized facies are

intergradational, and sharp, well-defined contacts between facies will generally not be observed:

1. A prodelta sequence, consisting primarily of clay and silt.

The prodelta can grade basinward into a continental shelf or conti-

nental slope, depending on the scale and geomorphic setting of the

delta. Precise distinction between prodelta and muddy shelf or slope

facies may be difficult or arbitrary.

2. An upward-shoaling, marine delta front sequence. Multiple

upward-shoaling (and typically upward-coarsening) cycles may be

superimposed on the overall upward-shoaling trend. Asquith (1970)

demonstrated a relationship between upward-coarsening delta-front

cycles and progradational clinoform geometry in the subsurface Upper

Cretaceous of Wyoming; this relationship is probably generally applic-

cable, but foresets are typically too low-angle to be perceived in the

field. Sedimentological details of the delta front depend strongly on

the physical regime of the marine basin and of the discharging river.

Galloway (1975) proposed a ternary classification for deltas with

river-dominated, wave-dominated, and tide-dominated end members; 283 differences among these end members should be strongly reflected in details of the delta-front facies. Delta-front facies can represent environments which also occur along non-deltaic coastlines (shoreface-

beach, strandplain, tidal inlet, tidal shoals, etc.), as well as uniquely deltaic environments (e.g. delta-mouth bar).

3. A sequence which is transitional upward into nonmarine

strata, and is characterized by environments of intermediate to low

salinity. The transitional unit may include deposits of lagoons,

brackish marshes, tidal flats, interdistributary bays, estuaries, and

tidal channels. In many deltaic sequences, particularly wave-dominated

ones, the transitional unit can be quite thin.

4. A delta plain sequence of nonmarine fluvial deposits,

including point bar, levee, and floodplain deposits; it may also

include thin, lacustrine, marsh, or swamp deposits.

In the FCV section, the succession of Mud Hills, Deguynos,

Camels Head, and Diablo members (Plate 8) can be readily interpreted in

terms of this generalized deltaic sequence, although boundaries between

genetic facies do not correspond exactly to the mappable units (Figure

6.10). Further interpretation of these generalized facies is based on comparison with other modern and ancient depositional systems, and with

studies of the modern Colorado delta (Figures 6.11, 6.12).

The FCV sequence is anomalous in some respects. For example, the delta-front cycles of the uppermost Mud Hills and lower Deguynos members do not culminate in typical shoreface-beach or delta-mouth bar facies (Coleman and Wright 1976, Coleman and Prior 1982); the capping

284

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Figure 6.11. Representative profiles from Meckel (1975) of modern Colorado delta, based on coring. Redrawn in same format as measured sections in Chapters 4 and 5. 286

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1!--1P9 1111---°9 1411— 287 sandstone and coquina bodies have more in common with models recently developed for shelf-sand bodies. In addition, rhythmite bedding of the type observed in the Mud Hills and Deguynos members is unusual for deltaic sequences. The transitional unit (upper Deguynos and Camels

Head members) is unusually thick (300-400 m) in proportion to the delta-front sequence (-200 m), compared to most deltaic sequences; this suggests that the delta front may have been considerably detached from the nonmarine part of the delta.

A possible modern example of a detached delta front is

suggested by recent work by Nittrouer et al. (1986) on the continental

shelf offshore from the Amazon River mouth. In that study, a zone of low-angle clinoforms with basinward progradation was observed in a

mid-shelf position. Unlike the classic conception of deltaic sedimenta-

tion, and the example illustrated by Asquith (1970), progradational

geometry of the Amazon deltaic system is not tied closely tied to a

shoreface or strandline. Nittrouer et al. (1986) have referred to this

situation as a "subaqueous delta".

Prodelta

The prodelta sequence in the FCV section consists of two and lithologies: massive, homogeneous, slightly silty claystone, sandstones. Lack rhythmite-bedded clayey siltstones to silty, very fine of intensive biotur- of fissility in mudrocks is probably an indicator

bation and therefore of a well-oxygenated water bottom (Byers 1974), in the although discrete burrows or burrow mottling were not observed apparent massive claystone facies. Intensive bioturbation is commonly 288 in the rhythmite facies, based on mottling, lack of distinct bedding contacts, and isolated burrows and burrow networks weathered out in relief.

Mode of deposition of the rhythmites and controls on cyclicity are problematic. While it is tempting to attribute the dominant cyclicity to annual seasonal-discharge cycles of the ancestral Colorado

River, it is not apparent how this model could be independently veri- fied. Recent studies of lacustrine varves (e. g. Lambert and Hs 1979) have indicated that such rhythmites may not be seasonal and may be more akin to turbidites, that is, episodic but not periodic.

Sedimentological evidence for the mode of rhythmite deposition is obscured by bioturbation; physical structures other than rare wisps of ripple lamination were not observed. The striking lateral continuity of rhythmite bedding is unusual for shallow-water, traction-dominated

deposits, but is quite common in fine-grained turbidites, which raises

the possibility that deposition could have been primarily by sediment

gravity flows.

Delta-front cycles

Individual upward-coarsening cycles in the upper Mud Hills for as much as member and lower Deguynos member can be traced laterally than that. As 6 km (Figure 6.13, Plate 7), and may be more extensive in the such they are the most correlatable architectural element

deltaic sequence, which is typical of delta-front facies in general.

Thickness variations in individual cycles may represent primarily of differential subsidence, although the dramatic eastward thickening

289

datum —mn m ft 100 300

FC-7 C-9

cross-stratification 100 rhythmite bedding (poorly exposed or developed) shells [coquina bed not exposed 0

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k A.iees channeled base

j Pdo Deguynos-Mud Hills contact 0'. Pm

Figure 6.13. Measured-section logs of delta-front cycles in upper Mud Hills and lower Deguynos members, simplified from Plate 7. 290 cycles in the Mud Hills member from Fish Creek Wash to the Elephant

Knees may be related in part to depositional paleotopography (Asquith

1970). The progradation direction of the delta is not known with any precision, but if it had an easterly component, the eastward thickening of these cycles could represent clinoform geometry. If that is the case, the most-extensive hogback-forming units, which define the base of the DeGuynos member, would correspond approximately with the change from topsets to foresets. The Mud Hills-Deguynos boundary is clearly

time-transgressive, becoming higher in the section in the direction of

apparent clinoform progradation (east). An easterly component of

progradation is compatible with fluvial paleocurrent directions in the

Diablo member (Figure 6.1) and with the inferred orientation of the

western basin margin, based on the geographic distribution of the 011a

member (Figures 1.6, 6.2, Plate 3).

Upward-coarsening cycles in the Deguynos member are typically

capped with sandstone or coquina beds; these vary considerably in

detail, which complicates development of a general sedimentological

model. Nevertheless, a few useful generalizations can be made. The

dominant structures are trough cross-stratification and larger-scale

accretionary stratification which may range from very low-angle to high

angle (-25°). The uppermost structures in sandstone and coquina beds, not where visible, are almost invariably cross stratification and

parallel lamination or ripple lamination. Paleocurrent directions the inferred from cross-stratification are strongly unimodal toward

south, although alternating north- and south-directed foresets occur 291 locally. Some potentially diagnostic features are apparently absent, such as reactivation surfaces (Klein 1977), "tidal bundles" (Kreise and

Moiola 1986), hummocky cross-stratification (Harms 1975), and graded

"tempestite" beds (Ager 1974).

High-energy, cycle-capping units of cross-stratified fine-grained sandstones and coquinas can be traced by correlating upward-coarsening cycles between closely spaced measured sections

(Figure 6.13, Plate 7). The most extensive high-energy unit documented

in this manner is the sandstone-coquina complex that forms the Elephant

Knees hogback. However, this unit consists of coalescing lenses that

are individually less extensive. The cross-stratified "notch sand" at

the base of the Deguynos between Fish Creek and North Fork is a resis-

tant ledge-former for little more than 1 km. Cross-stratified coquina

beds are typically continuous for 1 km or less, although they may come

and go at the tops of more continuous cycles. Narrow channeled sand-

stones with scoured bases, typically cut into rhythmites or bioturbated

siltstone or sandstone, occur locally below more continuous sandstone

bodies, as at western Elephant Knees (section EK-2) and Loop Wash

(section LP-1).

These cycle-capping sandstone and coquina bodies were clearly

high-energy, shallow marine deposits, but their precise mode of origin

is not readily apparent. Possible analogs for these sandstone and

coquina beds can be sought in the literature on coastal and

continental-shelf sedimentology, and a series of working hypotheses can

be considered. 292 1. Shoreface-beach model. The Deguynos cycles are superficially similar to cycles produced by progradational shoreface-foreshore

sequences (c. g. McCubbin 1982). However, such sequences typically culminate in parallel (planar) lamination or low-angle wedge cross- stratification of the foreshore, which overlies a trough cross- stratified shoreface facies. This sequence of sedimentary structures is not observed in the Deguynos member. Low-angle, accretionary foresets

in some coquina beds could presumably be attributed to a prograding

beach face. However, none of these beds, nor any of the Deguynos

delta-front cycles, is directly succeeded by indicators of subaerial

exposure; a very thick transitional unit intervenes between delta front

and delta plain, as previously discussed. Furthermore, tops of some

cycles have good marine indicators such as corals in the Elephant Knees

coquinas, including solitary Astrangia and tablets of ?Solenastrea

(S. M. Kidwell, 1986, oral communication). The lower Deguynos sandstone

and coquina bodies are more likely to represent marine shoal deposits

detached from the shoreline than shoreface-foreshore deposits.

2. Storm-dominated shelf-sand ridges. In recent years a number

of shelf sandstone bodies have been described from the Jurassic and

Cretaceous of the Western Interior Seaway of North America, including

the Cardium (Wright and Walker 1981, Krause and Nelson 1984), Sussex

(Berg 1975), Shannon (Davis 1976, Spearing 1976, Shurr 1984, Tillman and Martinsen 1984), Mosby (Rice 1984), Duffy Mountain (Boyles and

Scott 1982), and others reviewed by Boyles and Scott (1982), Walker

(1984), Swift and Rice (1984). Typical vertical sequences are 293 upward-coarsening from mudstone into sandstone and are suggestive of truncated strandline deposits, with shoreface-like sequences but without foreshore facies. Regional stratigraphic relationships indicate

that these high-energy sandstone bodies are detached from contempor-

aneous shorelines and are completely encased in shelf mudstones. Some

shelf sandstones, such as the Cardium and Mosby, are characterized by

hummocky cross-stratification (HCS) and graded beds, unlike typical

Deguynos sandstone bodies. Others, such as the Campanian Shannon,

Sussex, and Duffy Mountain Sands, are dominated by trough cross-

stratification with strongly unimodal paleocurrents, but HCS is rare or

absent, more like the Deguynos sandstones. Boyles and Scott (1982)

reported that paleocurrent directions in the Duffy Mountain range from

coast-parallel to obliquely onshore; paleocurrent data compiled by them

for other Upper Cretaceous sandstones range from obliquely offshore to

obliguely onshore, but all preferentially toward the south. The lower

part of the Shannon Sand is dominated by thinly interbedded sandstone

and mudstone, with abundant wavy and lenticular bedding, and prominent

burrow mottling (Davis 1976). In contrast, typical Deguynos sandstones

are characterized by rhythmite bedding, a bedding type that has not

been reported in association with shelf deposits of the Western

Interior.

The Deguynos also contrasts with Western Interior shelf sand- was probably stones in terms of paleogeographic setting. The Deguynos

deposited at the head of a long, narrow, high-relief, marine basin,

even narrower than the present Gulf of California (Chapter 8). Shelf 294 sandstones of the Western Interior were deposited in the outer part of a broad shelf facing a long seaway. It might be expected that the oceanographic regime of these two settings would be quite different.

3. Tidal-shoal or tidal-ridge model. Modern tide-dominated

deltas, including the modern Colorado delta (Thompson 1968, Meckel

1975) and Ord delta (Wright et al. 1975), are characterized by linear

shoals or sand ridges aligned subparallel to or radially from the river

mouth. Although such tidal ridges are quite common worldwide (Off

1963), they have proven difficult to study directly, and relatively

little is known about them compared to other shelf deposits. Further-

more, ancient analogs have not been convincingly demonstrated. Meckel

(1975) suggested that tidal ridges of the Colorado delta should produce

upward-coarsening cycles -10 m thick, with bimodal paleocurrent direc-

tions. Off (1963) indicated that such ridges range in amplitude from 3

to 30 m. Meckel (1975) demonstrated an upward-coarsening sequence

(Figure 6.11) for the tide-influenced shoreline sands at El Golfo on dip the eastern side of the delta, but they showed that dune foresets

toward the north.

Some examples of cross-stratified sandstones have been inter-

preted in terms of deposition on wide, tide-dominated shelves analogous

to the North Sea (Narayan 1971, Nio 1976, Johnson and Baldwin 1986). direc- Such sandstone sequences tend to have a dominant paleocurrent

tion, but also show features characteristic of tidal-current deposition

such as reactivation surfaces and herringbone cross-stratification. criteria However, according to Johnson and Baldwin (1986), diagnostic 295 for distinguishing between storm-dominated and tide-dominated shelf sands are not well established. Probably the strongest argument for tidal influence in the delta-front sequences of the Deguynos member is the tidal-flat characteristics of the overlying transitional facies.

Transitional (tidal flat) sequence

The thick transitional sequence (upper Deguynos member and

Camels Head member) consists of a wide variety of lithofacies and bedding types. The following features are thought to have particular genetic significance: (1) thinly interbedded ripple-laminated sandstone and claystone, referred to here as wavy bedding (sensu lato); (2) thick sandstone bodies with upward-fining or "blocky" (sharp base and top) profiles, similar to channel sandstones of the delta-plain sequence,

but with some notable differences; (3) a variety of shell beds and

shell concentrations; and (4) systematic up-section variations in

intensity of bioturbation, claystone color, and sandstone percentage.

Wavy bedding. Reineck and Wunderlich (1968) described a suite

of sedimentary structures from the North Sea tidal flats characterized

by thin interbedding of ripple-laminated fine sand and clay. Variation

in the percentage of sand and mud leads to a range of bedding types,

referred to as lenticular bedding (mud > sand), wavy bedding (mud

sand), and flaser bedding (sand > mud). Although this associataion of

bedding types was first described from tidal flats and is a common

feature of tidal-flat deposits (Klein 1977), the same association of

structures has been observed in thalwegs and point bars of tidal

estuaries (Howard and Frey 1975, Howard, Elders, and Heinbokel 1975), 296 in delta-mouth bar sequences of the Mississippi delta (Coleman and

Prior 1982), in upward-coarsening shelf sandstones such as the Creta- ceous Shannon Sandstone (Davis 1976, Spearing 1976), and in turbidites, where CE or CDE Bouma (1962) sequences can predominate locally, especially in Facies E of Mutti and Ricci-Lucchi (1978). Therefore, such wavy bedding (s. 1.) alone is not diagnostic of tidal-flat environments. Foraminiferal paleoecology of the transitional sequence in the

FCV section indicates a brackish-water to intertidal environment (Quinn and Cronin 1984), which is consistent with a tidal-flat interpreta-

tion. Wavy bedding in the FCV section does not occur in asymmetric

upward-coarsening cycles, as would be expected for shelf sandstones or

delta-mouth bars. In a few instances, it does occur in asymmetric

upward-fining sequences, which might indicate tidal estuaries. In

general, however, wavy bedding does not occur in well-defined asym-

metric cycles. Paleocurrents associated with the wavy bedding are the

most bimodal of any in the FCV section (Figure 6.1). Direct evidence

for reversing currents is difficult to observe, however, because ripple exposures lamination is generally obscure and difficult to measure, and sedi- of rippled bedding surfaces are rare. However, the aggregate of

entary features suggests deposition on channeled, non-vegetated, mixed

sand-mud tidal flats. gradual Channel sandstones. Within the Deguynos member, a

change from predominantly upward-coarsening sequences to upward-fining to aequenes effectively defines the change from delta-front 297 transitional facies (Figure 6.10). Exposures of the transitional sequence are generally inadequate to permit observation of lateral variations of the upward-coarsening sandstones, but they most likely represent channel deposits. They are similar to upward-fining sandstone bodies of the delta plain sequence (Diablo member), but are distin-

guished by common shell material, including basal shell lags, and by

wavy bedding, flaser bedding, and bioturbation, particularly in the

upper part of sandstone bodies. Trough cross-stratification is abundant

and paleocurrent directions are strongly unimodal, with less dispersion

than is typical for the Diablo member. This suggests that the sand-

dominated channels were mostly distibutary channels of the ancestral

Colorado River, which experienced some marine influence and which

reworked marginal-marine beds, but which were dominated by ebb-directed

currents.

Shell beds. A distinctive type of shell bed unique to the

transitional sequence consists of randomly oriented Ostrea shells (in

some cases with Anomia) in a silt or clay matrix. In such beds the

shells are commonly clumped or articulated. These beds probably repre-

sent in situ growth of oysters, with only minor reworking. Whether such occurrences beds represent the primary source for all the oyster-shell

in the deltaic Imperial Formation is unclear. In very thick, mud-matrix disoriented, oyster-shell beds of the upper Deguynos member, beds of

clumped shells alternate intergradationaly with beds of strongly

orisnted and presumably reworked shells. 298

Other shell beds and shell concentrations occur in the transitional unit that are neither obvious channel lags nor in situ oyster beds. These range from thin beds of oriented or disoriented oyster and/or pectinid shells, underlain and overlain by mudstone, to thick

(up to 2 m), well-bedded oyster coquinas with low-angle cross- stratification. These shell beds generally lack distinctive features that would permit a definitive sedimentologic interpretation. One possibility is that they represent strandline ridges or berms such as have been described from tidal flats of the modern Colorado delta

(Thompson 1968).

Up-section variation. Intensive bioturbation dominates the

lower part of the transitional unit (middle part of Deguynos member), so that bedding and lamination have been largely obliterated, and mottled mixtures of clay, silt, and sand are the dominant lithofacies.

Higher in the Deguynos member, some intervals show well-preserved

bedding and lamination, particularly wavy bedding, while other

intervals are thoroughly bioturbated. In the Camels Head member,

bioturbation is relatively sparse; wavy bedding and silt laminae are

typically well preserved. the Claystone color. The change of claystone color within

Camels Head member probably also indicates a distal-to-proximal facies (prior to change in the section. Suspended load of the Colorado River to dam construction) was distinctly reddish brown (Sykes 1937) due area. extensive fine-grained redbeds in the Colorado Plateau source

Sykes (1937) also reported that sediment deposited on the Colorado 299 River floodplain was light yellowish brown to dark chocolate brown, on the depending silt and sand content. As a general rule, iron is present in fluvial suspended load as hydrated oxides, and after deposition matures in situ to form red deposits (Collinson 1986). Sykes

(1937) reported that mud deposited in the estuary was chocolate brown to bluish gray, and attributed the change of sediment color in the estuary to reduction of ferric oxides to sulfides.

Reddish coloration in claystones of the Diablo, 011a, and

Camels Head members probably represents the original ferric-oxide

content of suspended load in the ancestral Colorado River. In the

Camels Head member, a gradation of reddish to gray colors is

observed. The gray color which is characteristic of the marine deltaic

claystones is probably secondary, due to reduction of the ferric

pigments as suggested by Sykes (1937). Reduction is most likely due to

incorporation of organic matter in the sediment, which would be favored

by subaqueous enironments. Therefore, the change of claystone color in

the Camels Head member probably reflects a transition from gray sub-

tidal to reddish supratidal facies. This color change occurs higher in

the section and therefore more proximally than the change from biogenic

to physical dominance of sedimentary structures (Figure 6.10).

Delta plain

The Diablo member is clearly differentiated into paleochannel

sandstones, represented by massive, upward-fining cycles, and overbank sequences of claystone and siltstone with thin to medium-bedded sand-

Stones. Sequence, and associations of sedimentary structures are mostly 300 compatible with well-established features of fine-grained meandering fluvial systems (e.g. Allen 1965a, 1965b, Collinson 1986), although some features are anomalous, as discussed below. The Diablo is the thickest and most extensively exposed unit in the western Salton Trough

(Plate 3). Volumetrically, this is probably the predominant unit filling the Salton Trough, and deposition has apprently been continuous to the present. Van de Kamp (1973) illustrated profiles of channel sands from the modern delta plain (Figure 6.12) that are very similar to profiles of the Diablo member.

Miall (1986) has recently criticized the practice of inferring paleogeomorphic aspects of fluvial systems such as channel pattern and sinuosity solely on the basis of vertical sequences. Preferably, the

architecture of channel deposits should be documented in two or even

three dimensions. Although recurrent upward-fining cycles associated

with abundant overbank muds are most characteristic of meandering

fluvial systems, similar sequences can be produced by braided streams

with high suspended load (Miall 1978). Interpretation of the ancestral

Colorado as a meandering river is therefore open to question, but

appears to be the most likely possibility because meandering distrib-

utaries of the modern Colorado delta plain have produced similar

vertical sequences (Figure 6.11, 6.12).

The greatest problem with a meandering-fluvial interpretation that would be is the paucity of well-defined lateral accretion surfaces concave expected from the lateral migration of point bars. Large-scale

Upward scour surfaces, defined by concentrations of mud-pebble 301 conglomerates, are well exposed along Fish Creek Wash above the Diablo turnoff, but these scour surfaces are not systematically organized, which results in a fairly random lensoidal geometry. Potter (1967) has referred to the geometry of coalescing lensoidal channel sandstones as "multilateral" sandstones; Campbell (1976) has documented such a fluvial system with geometry similar to that of the Diablo. Internal geometry of channel deposits in the Diablo may be dominated by vertical scour and subsequent backfilling during floods, as shown by Kniffen

(1932) for the Yuma gauging station (Figure 6.14), rather than by

systematic migration of point bars of constant cross-section.

Another anomaly in the Diablo member is the dominance of paral-

lel lamination over trough cross-stratification. Parallel lamination

typically occurs in the middle to upper point bar (Bernard et al.,

1970) but in most meandering fluvial channel deposits is a relatively

minor constituent (Collinson, 1986). However, Allen (1965b) illustrated

some upward-fining cycles from the Old Red Sandstone that are dominated

by parallel lamination, similar to the Diablo member. Van de Kamp

(1973) demonstrated that upward-fining cycles of the modern Colorado

delta plain are also dominated by parallel lamination (Figure 6.12).

The prevalence of parallel lamination in Colorado River paleochannels may be largely a function of grain size. Flume experiments indicate that formation of dunes (i.e. medium-scale ripples) is suppressed in very fine sand, and plane beds and small-scale ripples predominate

(Middleton and Southard, 1977, p. 7.37). This result is also consistent

With the observation that cross-stratification in Colorado River 302

100 ZOO 4P00 400 500 000

- 8 AR- \1200

no° CFS.

0 3

, 1921 Maximum. ertiSs Section Colorado River June£8 • • Sept 11, 192.4

Figure 6.14. Profiles of Colorado River channel at Yuma gauging station, at maximum and minimum discharge during a 24-year period before construction of Hoover Dam. From Kniffen (1932). 303 paleochannels is commonest where slightly coarser "L"-suite sand is admixed.

Convolute lamination is particularly characteristic of paleochannel sandstones in the Diablo member. Some medium-scale convolutions are clearly the result of oversteepened foresets in cross- stratification, but most convolute lamination in the Diablo paleochannels is too large to be readily explained this way. Similar

large-scale convolutions have been observed in point-bar sequences along the lower Mississippi River (Coleman and Prior 1982), and can be

readily explained as a waning-flood phenomenon. A rapid fall of stage

results in a large head gradient between pore water in the point bar

sand and water in the adjacent channel. The head gradient can result in

rapid dewatering and consequent liquefaction and soft-sediment

deformation of the point-bar sand.

Paleocurrent data from the Diablo member and "C"-suite sand-

stones of the 011a member clearly indicate that the FCV section was

located on the southeast, or Gulfward, flank of the Colorado delta with plain during the period 4.0 to 2.8 Ma. These data are incompatible

the present position of the FCV section to the west of the modern delta If, plan and entry point of the Colorado River into the Salton Trough. of as seems likely, the ancestral Colorado entered the Trough by way data the same gap between the Chocolate and Laguna Mountains, these

imply tectonic displacement to the northwest since the Diablo member

was deposited (Winker and Kidwell 1986). 304 Paleoecological problems in the deltaic sequence

Paleoenvironmental interpretations of the deltaic sequence based on micropaleontology (Figure 6.10, Quinn and Cronin 1984)) are

generally consistent with those based on sedimentologic evidence, at

least as far as relative water depths and salinity. Terms such as

"inner shelf" and "outer shelf" are somwhat misleading and inappropri-

ate in this depositional setting

Paleoecological interpretation of the marine macrofaunas is

much more problematic. The dominant Ostrea-Anomia-assemblage occurs

throughout the deltaic sequence (except in the prodelta) despite

obvious sedimentologic differences. Some faunal constituents, such as

Fasciolaria, Balanus, and encrusting bryozoans, are actually most

typical of the transitional Camels Head member, and a diverse fauna was

found in the laterally adjacent Jackson Fork member. The occurence of

fairly diverse marine assemblages in what should have been stressful,

brackish-water environments is rather puzzling.

The nature of the fossil occurrences is at least as important

as the fossils themselves. Oyster-shell beds with muddy matrix are the

most likely to be in situ accumulations; shells in other occurrences

may have been significantly transported. Coquina beds of the delta-

plain sequence may have accumulated from offshore transport of shells are reworked from in situ oyster beds; offshore-directed paleocurrents

consistent with this model. Some of the shell material in coquina beds in is in situ, however, based on the occurrence of articulated Anomia

Deguym os coquina beds in the Coyote Mountains. 305 Accumulations of shell material higher in the section require a different explanation, particularly in channel sandstones of the upper Camels Head member and nonmarine Diablo member, where cross-

stratification does not indicate onshore transport. Two modes of

reworking are envisioned: (1) short-distance reworking by lateral

channel-cutting into fossiliferous beds, and (2) longer-distance

reworking from outcrops of uplifted fossiliferous strata. The latter

mechanism is required to explain the occurrence of marine shells and coquina clasts in the Diablo member.

Summary of the deltaic sequence

The deltaic sequence consists of a laterally persistent

succession of stratigraphic units (Mud Hills, Deguynos, Camels Head,

Diablo). Individual beds, depositional cycles, or sandstone/coquina

beds, are characterized by lack of lateral continuity. An exception is

the delta-front cycles, which are continuous between Fish Creek Wash

and North Fork and possibly farther (Figure 6.13). This is a common

observation in deltaic sequences, that upward-coarsening cycles within

the marine part of the delta are the most laterally persistent strati-

graphic element. As discussed previouly, an anomalous aspect of this

sequence is the considerable thickness of the transitional zone between

delta front and delta plain.

Evidence for tidal influence on deposition comes more from

sedimentary structures within the transitional zone than from paleocurrent data. The distinctive wavy bedding (s. I.), though not in

itself diagnostic of tidal influence, in association with 306 paleoecological evidence and stratigraphic position indicates tidal-flat environments. The most proximal (stratigraphically highest) tidal-flat strata are very similar to the delta-plain strata except for (1) the wavy bedding which appears in what would otherwise be inter-

preted as floodplain deposits, and (2) the common concentrations of shell material in channel sandstones. Lower and more distally in the

tidal-flat sequence, occurrences of gray claystone, in situ oyster

beds, and thin coquina beds indicate a stronger marine influence, and

probably intertidal to subtidal environments. In the lowest, and pre-

sumably most distal part of the transitional zone, bioturbation and

biogenic sedimentary structures dominate over physical structures.

The origin of high-energy shoal deposits capping delta-front

cycles is more ambiguous. These sandstone and coquina beds resemble

deposits that have been interpreted as storm-dominated shelf deposits

in the Cretaceous and Jurassic of the Western interior, although the

geographic setting almost certainly differed in terms of basin shape,

tectonic setting, and the nature of contemporaneous shorelines.

Evidence for tidal influence in the transitional zone would tend to

support a tidal-shoal interpretation, but unimodal paleocurrents and

lack of reactivation surfaces or "tidal bundles" would tend to argue

against this interpretation. At the present time, diagnostic criteria for tide-dominated versus storm-dominated shelf sands have not been developed (Johnson and Baldwin 1986).

Channel sandstones with scoured bases, lenticular geometry, and ulAmTd-fining to blocky profiles occur throughout the deltaic sequence, 307 except in the prodelta. These channel sandstones are most abundant in the delta-plain and proximal tidal-flat settings, and become less common down-section. Paleocurrent directions in these channel sandstones are consistently unimodal and south-directed (or ebb- dominated). This suggests that these were fluvial and distributary channels rather than tidal channels. Deposits of tidal channels may occur within the sequence, but these could be sand-poor and difficult to distinguish from tidal-flat deposits, given the limited lateral

extent of outcrops and the difficulty of seeing bedding geometry. In the ternary classification of Galloway (1975), the Pliocene Colorado would probably be placed between the river-dominated and tide-dominated poles, depending on how the evidence for and against tidal influence is weighed.

In comparison with models for the modern Colorado delta, the

FCV section is generally similar, but different in some details.

Sedimentary structures and sequences of the delta plain, and its association with lacustrine deposits, closely resemble those predicted by Van de Kamp (1973). Extensive tidal flats appear to be common to the modern and Pliocene Colorado deltas. However, evaporite beds and evaporite-disrupted chaotic mudstones, documented by Thompson (1968) for tidal flats of the modern delta, were not observed in the FCV section. Linear tidal shoals, discussed by Meckel (1975), may have an ancient counterpart in sandstone bodies of the Deguynos member.

DPward-coarsening delta-front cycles and rhythmite bedding, two of the 308 most striking features of the marine deltaic sequence, are not predicted by the Meckel (1975) model.

Lacustrine units

Lacustrine units in the Salton Trough are distinguished from marine units by the lack of open-marine fauna and paucity of bioturba-

tion, and are distinguished from fluvial units by the thin, laterally

persistent bedding geometry and predominance of mudstones over sandstones. These lacustrine units can be divided into those dominated by

"C"-suite and "L"-suite sandstones and siltstones. The stratigraphie

setting of lacustrine units in the Vallecito Basin and Borrego Badlands

(Plate 3) is generally similar despite differences in nomenclature (Figure 5.2).

In the Vallecito Basin, the Tapiado member contains exclusively

"L"-suite sandstones and siltstone beds, although the provenance of clays is not known. The Tapiado rests with sharp contact on the Diablo and 011a members; this contact marks the uppermost occurrence of "C"- suite sandstones. The Tapiado grades laterally and upward into braided- stream sandstones of the Hueso member. This zone of intergradation is marked by thin interbedding of claystone, siltstone, and sandstone, with an abundance of ripple lamination (especially ripple-drift or climbing7ipple lamination), and centimeter-scale to decimeter-scale convolute lamination. In Tapiado Wash, this zone of intergradation contains upward-coarsening cycles on the order of 10 m thick. This association of structures is typical of small lacustrine deltas.

Large-scale foreset ("Gilbert delta") geometry was not observed, 309 however, suggesting that these lacustrine deltas had low-gradient slopes.

In contrast to the Tapiado member, the Borrego and Brawley

Formations are lacustrine units dominated by "C"-suite sandstones and

siltstones. The Diablo member grades upward into the Borrego Formation;

in this gradational zone, thick channel sandstone and red overbank

mudstones typical of the Diablo member alternate with sequences of gray

claystone and pale-orange siltstones typical of the Borrego. This

gradational zone with alternating fluvial and lacustrine sequences

probably represents rapid fluctuations of the lake level, similar to

the Holocene Salton Sea-Lake Cahuilla. Rapid lake-level fluctuations

controlled by discharge of the ancestral Colorado River could explain

the general absence of upward-coarsening lacustrine-delta cycles in the

Borrego Formation.

"C"-suite sandstones in the upper Borrego and Brawley Forma-

tions in the San Felipe Hills have predominantly west-directed paleocurrents (Wagoner 1977). These sandstones were probably reworked from

older "C"-suite units that were uplifted on the westernmost side of the

Salton Trough. The age of these sandstones, which have generally been basin placed in the Pleistocene, may mark the initial uplift of Neogene initial fill in the western Salton Trough, and could thus be related to

(or renewed) movement on the San Jacinto fault system.

Evaporite beds, which might be expected in lacustrine sequences

in an arid setting, were not observed in the Tapiado member or in the

Borrego Formation on the western side of the Salton Trough. Babcock 310

(1974) reported gypsum beds from the Borrego Formation on the northeastern side of the Trough, adjacent to the Salton Sea. This probably means that nonmarine evaporite deposition was confined, for the most part, to the topographically lowest part of the lacustrine basin (as in the modern Salton Trough), and not immediately adjacent to the ancestral Colorado delta plain.

Compared with the widespread Borrego Formation and its grada-

tional contact with the Diablo member, the Tapiado member and its sharp

contact with the Diablo member appear to be local and anomalous (Figure

5.2). A possible explanation for the abrupt cutoff of Colorado River

sediment is that the topographic low between the Colorado delta plain

and basin-margin alluvial fans shifted toward the east, and a lake

occupied this topographic low. This explanation is unsatisfying,

hlowever, because comparable lacustrine beds do not occur in the

mixed-provenance zone (011a member) lower in the section. An alter-

native explanation is that a topographic barrier developed between the

Colorado delta plain and the FCV area, thereby excluding Colorado River are sand and silt. Elongate ridges with outcrops of Neogene strata

commonly uplifted along active strike-slip faults; Crowell and

Sylvester (1979b) have referred to such uplifts of formerly subsiding

blocks as "porpoise structures". Uplifted porpoise structures are

omm in the Salton Trough region; one example is the Sierra de los

Cucapas along the Laguna Salada fault. This range separates Laguna

Salada from the Colorado delta plain, and may be analogous to the 311 situation that resulted in persistent "L"-suite lacustrine deposition in the FCV section.

Basin-margin alluvium

Terrigenous sediment was shed from the western flank of the

Salton Trough continuously during deposition of the Colorado delta,

from the early Pliocene to the present. While it is useful for the purposes of mapping and description to subdivide these locally-derived

strata into various units (Canebrake, 011a, Jackson Fork, Hueso,

Vallecito, Ocotillo), it is important to realize that deposition of

basin-margin alluvium was continuous and that placement of strati-

graphic boundaries and application of nomenclature are somewhat

arbitrary. Waxing and waning of basin-margin alluvial units can

represent responses to a variety of phenomena, including pulses of

uplift of the basin margin, reorganization of drainage systems along

the basin margin, and independent controls on the extent of delta

plain, marine, or lacustrine sedimentation, as previously discussed.

Individual controls on such pulses of basin-margin sedimentation are

generally difficult to isolate.

Margin of delta plain

The 011a member is of special importance in that it represents

the interfingering of locally derived alluvium of an extensive basin is margin bajada, with deposits of the Colorado delta plain, and of the delta therefore an unmistakable indicator of the western margin of plain. "C"-suite intervals in the 011a are virtually identical to 312 the Diablo member, except that concentrations of the "L"-suite sand and gravel typically occur in the lower portions of upward-fining paleochannel sandstone sequences, and trough cross-stratification is correspondingly more common. These channel lags of coarse "L"-suite

sediment are attributed to local input of locally derived sediment to

the Colorado River where it impinged on the basin-margin bajada.

"L"-suite sequences are more disorganized than "C"-suite

sequences, and overbank mudstones are much less abundant. For these

reasons a braided-stream interpretation is preferred to a meandering-

stream interpretation. Medium to large-scale tangential cross-

stratification is common in part of the 011a member, and probably

represents transverse bars; the foresets provide the best paleocurrent

indicator in the locally-derived part of the 011a member. Olive to

greenish mud drapes and mud cracks are common in these sequences, and

probably represent desiccation of stream beds between flood events.

Mudlstone intraclasts are typically thin and tabular, and probably

represent reworking of desiccated mud drapes. Dark-colored debris-flow beds also occur in the 011a, generally higher in the section and presumably more proximally than the predominantly cross-stratified part.

However, debris-flow beds are much less common than in the Elephant

Trees fanglomerates. Upward-coarsening and upward-thickening "L"-suite sequences, as observed in Layer Cake Wash, may indicate progradation of the basin-margin bajada during periods when the Colorado River migrated away from the basin margin. 313

Outcrops of the 011a member occur in the Santa Rosa Mountains, western Borrego Badlands, Borrego Mountain, FCV section, and the vici-

nity of Sweeney Pass (Figure 6.15; Plate 3). In all cases the 011a is

part of a gradational facies tract between the Canebrake Conglomerate

and Diablo member, except in the Borrego Badlands where the Canebrake is not exposed. These localities can be used to map the western margin

of the delta plain; in practice it is convenient to map both the eas-

tern limit of Canebrake Conglomerate and the eastern limit of Diablo

member (Figure 6.15). On a present-day (nonpalinspastic) base map,

these lines define an approximately north-south trend, with a large

recess west of Borrego Mountain (Figure 6.15).

This trend has probably been distorted by strike-slip tectonics

associated with the San Jacinto and Elsinore fault systems. For a

first-approximation palinspastic reconstruction, known tectonic

displacements can be restored and the limits of Canebrake and Diablo

distribution replotted (Figure 6.16). 25 km of right slip was restored

on the San Jacinto fault system (Sharp 1967) between the Santa Rosa and

Borrego Mountain localities. Apportionment of total San Jacinto offset the Borrego between the Clark and Coyote Creek faults is not known, so After this Badlands locality was not included in the reconstruction. of Borrego restoration is made (Figure 6.16), the sharp recess west

Mountain on the present-day base (Figure 6.15) is eliminated. on 3 km of right slip was restored on the Elsinore fault, based

the apparent offset of the Canebrake-011a-Diablo facies tract. The FCV

locality was retro-rotated to correct for -35° of post-depositional CW

314

9 10 15 20 25 km

4., 1). •G< • )‘ • ----.—....„„\ 80.R ******* * + • 15'1 ,, „,1 Po BADLANDS . • •Z ,, 0 -.2r) . Pd • t:k.I (3-Pd? .i.9- . ss,c3 Pd C:I. • li • . <,,,,, ri 5 . D Cr.. • '0,* (gPo -- CD o , • Po 17% SAN s . BORREGO • FELIPE . MTN . • • % HILLS .B.— -4 .±. •• • • • i cr"" +%• • 4, + • c4o • .....-.: a) a) '•...... • • ...E to` c7 •• • • ...... • c .0 C D La) E a) c 0 . • og° g • .(7)..6(3 • • FISH CRK. 334_ BASIN + ss-- ••• 00'

assumed effective diameter of rotated block = 20 km

3221_ 45' 1164T5' 110+60'

and Canebrake Figure 6.15. Distribution of Diablo and 011a members western Salton Trough, simplified from Plate 3. Eastern Conglomerate in the trend of limit of Canebrake and western limit of Diablo approximate the western margin of the delta plain. 315

0 5 10 15 20 25 km

'ball bearing model for 35 0 rotation of FCV section requires 12 km of shear

ALTERNATIVE 2

POc

Pd ALTERNATIVE 1

Figure 6.16. Suggested palinspastic reconstruction of Figure 6.15. 316 rotation (Johnson et al. 1983). The amount of regional shear required to accommodate this much rotation by "ball-bearing" tectonics depends on the effective diameter of the rotated block. A diameter of 20 km was assumed, which requires -12 km of distributed right slip. Johnson et al. (1983) apparently assumed a smaller effective diameter, and estimated only 7 km of right slip to account for block rotation. Relative motions between the rotating FCV block and the Sweeney Pass and Borrego

Mountain localities are not known. In the simple "ball-bearing" restoration, these two area become considerably separated ("Alternative 1" in

Figure 6.16). Alternatively, the Sweeney Pass locality may have remained coupled to the FCV block during rotation ("Alternative 2" in

Figure 6.16). The correct restoration is likely to be somewhere between these alternatives.

Mean paleocurrent directions for Colorado River paleochannels in the FCV section are approximately south-southeast in the present geographic orientation, but southeast to east-southeast when corrected for block rotation (Figures 5.1, 6.1). The corrected directions are somewhat divergent from the inferred margin of the delta plain in

Alternative 1, but are essentially parallel to the inferred margin of the delta plain in the FCV area in Alternative 2. Paleocurrent data from "L"-suite sandstones of the 011a member indicate dominantly eastward transPort, with a high degree of dispersion (Figure 5.15). These could be interpreted as a combination of transport perpendicular and parallel to the basin margin. 317 The resulting picture is one of meandering channels of the Colorado delta plain flowing on average parallel to the basin margin (i.e. south to southeast. Alluvial fans and braided streams shed

locally-derived sediment away from the basin margin, but braided

streams were deflected southward, between the basin margin and the delta plain.

Basin-margin fanglomerates

The Canebrake is typically interpreted as an alluvial fan

deposit. This is probably a reasonable approximation, but the Canebrake

is much different in appearance from the Elephant Trees fanglom-

erates. In particular, debris-flow beds are quite rare in the

Canebrake, and bedding contacts are not sharply defined. Crudely bedded

sand-gravel strata similar to that of the Canebrake can be seen in

erosional cuts into low terraces along the modern washes of the Fish

Creek drainage. Thus, the Canebrake may be more like a gravelly

braided-stream system in which debris-flows were not a major

process. Still, the rapid decrease in clast size away from the basin

margin, on the scale of a few kilometers, is more consistent with

alluvial-fan models than with gravelly braid plains (Rust and Koster 1984).

Basin-margin tidal flats

The Jackson Fork member, though small in extent, is important

for its diverse and previously undescribed macroinvertebrate fauna. The close association of full marine macrofossils and camel hoofprints 318 indicates very shallow-water or intertidal environments; inferred lateral continuity with the brackish-water to intertidal Camels Head member supports this interpretation. The association of coarse sediment, rapid lateral facies changes, and debris-flow beds is suggestive of a small alluvial fan constructed on the edge of the Colorado delta tidal flats.

The presence of diverse marine fauna in a tidal-flat setting adjacent to the ancestral Colorado delta presents a paradox, especially when compared to the relatively impoverished fauna of the Deguynos and

Camels Head members. A partial explanation may be that if the Colorado

River discharged temporarily into an ancient counterpart of the the

Salton Sea, the head of the ancestral Gulf could have reverted during

that time to full marine salinity, long enough for a diverse marine

community to establish itself. However, this hypothesis alone does not

explain why similar evidence for alternating brackish and full-marine

conditions does not appear in the macrofauna or microfauna of the

Deguynos and Camels Head members.

Summary

The overall stratigraphic succession of the FCV section is

consistent with what might be expected in a young rift to pull-apart of the basin filled mainly by a large fine-grained delta. The beginning

Salton Trough or proto-Gulf tectonic province began with early to

middle Miocene rifting, which disrupted pre-existing regional drainage,

and caused a change from deposition by through-flowing bedload streams and (Red Rock member) to local alluvial fans (Elephant Trees member) 319 basalt flows (Alverson Formation). Extensional subsidence led eventually to marine transgression of rugged, high-relief topography. Immediately prior to marine inundation, evaporite beds of limited extent but exceptional purity were deposited locally, perhaps in a "deep-basin, shallow-water" setting. Early marine deposits were diverse and laterally quite variable, but were generally coarse-grained and associated with alluvial fans and fan deltas. Shallow-water fan fringes and rocky outcrops were occupied by a typical Gulf of California marine macrofauna. Due to low wave and tide energy, sandy bottoms were

thoroughly bioturbated, although coarse clasts were episodically

reworked during flash floods or storms. Deeper-marine fan fringes were

dominated by high-density turbidity currents on high-gradient submarine

slopes. Two catastrophic failures of active fault scarps created megabreccia boulder beds, one nonmarine (subaerial?) and the other submarine.

Input of Colorado River sediment into the Salton Trough caused a major change of sedimentary patterns in the early Pliocene, as large volumes of clay, silt, and fine sand rapidly overwhelmed coarser- grained locally-derived sediment and displaced full-marine conditions farther to the south. The first Colorado River sediment was introduced as a small submarine fan of relatively low gradient, which produced a

Classical turbidite beds. As the basin was rapidly filled in, turbidites were succeeded by prodelta clays and silts and then cycles of shallow-marine, delta-front progradation. Most delta-front cycles culminated in higher-energy deposition of sand and shells, possibly 320 tide- or storm-dominated. These were succeeded, with continued progradation, by extensive channeled tidal flats. Finally, nonmarine environments prevailed, as a delta plain with meandering fluvial channels prograded southward. During all this time, local alluvial fans and braided streams fed coarse arkosic sediment from the western basin margin. This locally derived sediment interfingered along the basin margin with the volumetrically dominant fine-grained sediment from the

Colorado River.

Supply of Colorado River sediment to the FCV section was abruptly terminated at 2.8 Ma, possibly due to creation of a topographic barrier such as an uplifted "porpoise structure". Subsequent

sedimentation in this area was dominated by locally-derived lacustrine, braided-stream, and alluvial fan deposition. During the last million

years, the section was uplifted, tilted to the SW, and rotated CW by

-35° (Johnson et al. 1983).

The onset of strike-slip tectonism in the western Salton Trough

is difficult to date precisely. The Canebrake-011a-Diablo facies tract,

which indicates intergradation of locally-derived alluvium with

Colorado delta-plain alluvium, continues unchanged across the San

Jacinto and Elsinore faults; there is no evidence of local sediment

being shed from along these faults during the early or middle

Pliocene. On the other hand, this does not rule out tectonic activity; with the active Imperial Fault crosses the modern Colorado delta plain

little or no topographic expression and presumably little sedimentary

manifestation. 321 Termination of the Colorado River sediment source in the FCV area at 2.8 Ma may be the earliest manifestation of strike-slip sedimentation in the western Salton Trough. Eastward reworking of

Colorado-derived sediment into the San Felipe Hills, inferred from evidence presented by Wagoner (1977), may also reflect incipient tectonism and uplift in the western Salton Trough. The timing of this episode has not been closely constrained, but is probably early

Pleistocene. Uplift, tilting, and rotation of the FCV section, now constrained to the last 0.9 Ma (Johnson et al., 1983), is the clearest indicator of the onset of transcurrent tectonism in the western Salton

Trough. CHAPTER 7

REGIONAL STRATIGRAPHIC AND CHRONOLOGIC SYNTHESIS

Conventional stratigraphic nomenclature for the Neogene of the

Salton Trough developed largely through a series of historical acci- dents, and is only coincidentally related to genetically significant rock units. For this reason it was decided to recast the regional

Neogene stratigraphy in terns of informal genetic-stratigraphic units as discussed in the previous chapter, using the FCV locality as a reference section (Figure 7.1). Each of these units is physically distinctive and recognizable in the field, and could therefore be formally defined as a stratigraphic unit. However, it is unlikely that

completely revised nomenclature would be widely adopted, so all of the units are informal and have been given numerical and interpretive

designations. Relationships between these informal units and formal

stratigraphic nomenclature are summarized in a series of synonymy

charts (Figures 1.8-1.11, 7.2-7.8). Genetic stratigraphy of the FCV,

CIA, and Coyote Mountains sections was summarized in the previous

chapter and are not be repeated here.

322 323

o 0. "En E 010 0 0 76 0 = '18 OU 10' 15 (6) CO < 12 Z .0 D CO .0 F2P.1 Z ta C.- 0 .0 ce V) .E - m 1 •,1' •11' MA/Wœ v Z cJ co CC CO < 0 < • u- F- CO LI- 03 C9 o •

••••

"c3. cc)

o o

-o o

o 'CZ .w 0 9--, 0

Brooks & Roberts, 1954 Miller,1935 this study Hawkins, 1970 Minch & Abbott,1973

Jacumba Jacumbe 2 Basalt Basalt

Table Table Mountain Mountain* 1 Gravels Gravels

*first use of name

Figure 7.2. Stratigraphic synonymy for Jacumba area. 325

.Dibblee, Barnard, 1968 this study 1954

Canebrake 7 Formation

Palm Palm Spring Spring Formation 5 Formation

Imperial Imperial 4 Formation Formation

3 Split fanglomerates" Mountain ColoMa 2 Formation Progreso Volcanics

*first use of name

Figure 7.3. Stratigraphic synonymy for Sierra de los Cucapas 326

Vaughan, Bramkamp, Allen, Proctor, Peterson, Powell, this study 1922 1935 1957 1968 1975 1986

Painted Painted Painted Painted Indio' Hill* Hill Hill Hill 7 formation Formation Formation Formation Formation

wd Burrobend 70 member Lion* Imperial Imperial Imperial Imperial 3a sandstone formation Formation Formation Formation T3 Latrania E member

Hathaway Coachella* Coachella Coachella Coachella Hathaway Formation conglom- Fanglom- Fanglom- Fanglom- 2 erate formation Coachella erate erate erate fanglom.

*first use of name

Figure 7.4. Stratigraphie synonymy for vicinity of Whitewater Canyon (Super Creek and Lion Canyon). 327

Buwalda & i D bblee, Stotts, Proctor, Powell Free, 1914 Stanton, ' this study 1954 1965 1968 1986 1930

a) ._ Indio as Hills ,_ En 0 E, c 2 .cc g Formation ._ 0 Indio* c 0.= w i-, = co co co 2 w E 7 Mud Hill* formation ° E E 8 To o Thousand 2 Tz u_ a_ I'L Palms* al Q. Series Formation 0

Willis . IL Burro- Imperial Imperial - Carrizo Palms 2 bend 4c, 4b formation Formation Formation t; Formation a mbr. E

*first use of name Willow 1Stotts places this Hole*Fm.1 below Willis Palms, probably incorrectly

Figure 7.5. Stratigraphic synonymy for Indio Hills. 328

Dibblee, Babcock, 1969, this study 1954 1974

Borrego & Borrego Brawley 6 Formation Formations

Skeleton L .. Canyon* *— Palm Spring 76 Member

Formation co 7 upper -43> Sea .ce View* (i) Mbr. lower

*first use of name

Figure 7.6. Stratigraphic synonymy for Durmid Hill. 329

Olmsted Winterer Eberly & et al., . Stanley, this study 1975 1973 1978

Pliocene & older younger alluvium Colorado 5 (old Colo. & sand & Gila R.alluv.) gravel

'transition 'transitional 4d zone' unit

Bouse Bouse Bouse 4(b?) Formation Formation Formation

'Marine 'older marine 'Older . upper sedimentary 4a marine Miocene rocks' unit strata"

'nonmarine Tertiary sedimentary & 2 redbeds rocks' sediments"

Figure 7.7. Stratigraphic synonymy for Fortuna and San Luis Basins. 330

Metzger, 1968; Smith, 1970;, Metzger Wilson, 1931. Hamilton, et al., 1973; Eberly & Buising,1986 this study 1933,1962 1960 Metzger & Stanley,1978 Loeltz,1973; Winterer, 1975

o Colorado • 3 (0 River os ,,-,' Tu '6 > cl 8 alluvium '5 cr ,- 0 °'

unconsolidated * inter- clay,, silt, oc_ bedded 4 sand and au unit gravel E 11 Bouse Bouse *Cibola '6O u_ beds* W m.2 Fm. Formation cf, basal z 'lime cape „9, 3a — slitmoene

fanglom- Osborne Wash*ash* 2

*first use of name

Figure 7.8. Stratigraphic synonymy for Bouse Embayment (type area and Milpitas Wash). 331 Genetic-stratigraphic units of Salton Trough

Unit 1. Pre-rift nonmarine clastics

This unit is represented in the FCV section by the Red Rock

member area, which is interpreted as a predominantly pre-rift unit because of clast composition and paleocurrent directions which suggest

extra-local derivation (Kerr 1982, 1984) and therefore regionally

integrated drainage. The lower age of this unit is essentially uncon-

strained, except that it post-dates plutonic basement. The Eocene(?)

Ballena gravels in the Vallecito Mountains (Kerr 1982) could arguably

be included in Unit 1. The Ballena Gravels occupy north-northwest

trending paleovalleys, which is consistent with paleocurrent directions

for the Red Rock member, but the Ballena contains a much higher per-

centage of extra-local clasts than the Red Rock (Kerr 1982). The Table

Mountain Gravels in the Jacumba area contain at least 50% extra-local

clasts and discordantly underlie the Miocene Jacumba Volcanics (Minch

and Abbott 1973); the Table Mountain is therefore also regarded as a pre-rift unit.

Unit 2. Syn-rift nonmarine rocks.

This unit comprises fanglomerates (Elephant Trees member and

Coachella Fanglomerate), a nonmarine megabreccia (Lower Boulder Bed),

and alkalic to tholeiitic basalts (Alverson and Jacumba), all of which

underlie the lowest marine deposits, and therefore pre-date the proto-Gulf transgression. Apparently syndepositional normal faults 332 occur in Split Mountain (No Return fault) and at Whitewater Canyon (Peterson 1975).

Unit 3. Pre-deltaic marine units

Marine facies of exclusively "L"-suite sandstones (arkose and volcanic litharenite) and associated locally-derived conglomerates constitute this unit. It can be further subdivided into Subunit 3a, shallow-marine facies with abundant megafauna, (Latrania member, proximal Stone Wash member, and Imperial Formation at Whitewater

Canyon), and Subunit 3b, deeper-water units dominated by gravity-flow deposition, (Lycium, distal Stone Wash, and Kamia Creek members and Upper Boulder Bed).

Unit 4. Deltaic marine units

These units consist primarily of "C"-suite sandstones and associated siltstones and claystones, and contain marine macrofauna and microfauna. This unit is subdivided into four subunits, based on members of the FCV section: Turbidite sandstones like the Wind Caves member (Subunit 4a), prodelta claystone and siltstone like the Mud

Hills member (Subunit 4b), delta-front cycles with coquina beds, like the lower Deguynos member (Subunit 4c), and a transitional facies like the upper Deguynos and Camels Head members (Subunit 4d).

Unit 5. Delta plain

This unit consists of reddish, predominantly fine-grained fluvial sediments with "C"-suite sandstones (Diablo member and

"C"-suite facies of 011a member). 333

Unit 6. Lacustrine deposits

This unit consists of a variety of units, predominantly gray to olive claystones with freshwater and/or brackish-water faunas,

associated with siltstones and sandstones of both the "C"-suite (Borrego and Brawley Formations) and "L" suite (Tapiado member; Borrego

Formation in Durmid Hill).

Unit 7. Basin-margin clastics

These strata consist of "L"-suite sandstones and associated

conglomerates (e.g. Canebrake and Ocotillo Conglomerate, Hueso and

Vallecito members), and include mixed-provenance units (011a and

Jackson Fork members).

Unit 8. Colorado terrace deposits

This unit is limited to the Bouse Embayment, where Colorado

River alluvium is entrenched in the the Bouse Formation (Metzger et al.

1973, Metzger and Loeltz 1973), rather than conformably overlying the

Bouse, as in the Yuma Desert (Olmsted et al. 1973).

Local Summaries

San Felipe Hills, Borrego Badlands, and S. Santa Rosa Mountains (Figure 1.11) at In the San Felipe Hills, the lowest exposed unit crops out

Shell Reef (Plate 3) and is equivalent to the Deguynos member (Subunit Diablo 4c); it grades up through a transitional sequence (4d) into the

equivalent (Unit 5). The Diablo equivalent grades up into the

lactmtrine Borrego Formation (Unit 6). Another lacustrine unit, the 334 Brawley Formation (also Unit 6) is separated from the Borrego Formation by a tongue of the Ocotillo Conglomerate (Wagoner 1977, Dibblee 1984). The Ocotillo and Canebrake Conglomerates are both regarded as basin-margin elastics (Unit 7). The Canebrake grades into a mixed-provenance 011a equivalent (Canebrake facies B of Hoover 1965) which in turn grades into the Diablo equivalent. Sandstone and conglomerate below the Imperial Formation have been reported from wells in the area and assigned to the Split Mountain Formation (Dibblee 1984); these may represent the syn-rift nonmarine sequence (Unit 2).

Yuha Basin

A partial sequence is exposed in the Yuha Basin east of the

Elsinore Fault: a deltaic-marine unit with prominent coquina beds, exposed at Yuha Buttes (Plate 3), equivalent to the Deguynos member

(Subunit 4c), a transitional unit equivalent to the Camels Head member

(Subunit 4d), and a delta-plain sequence, equivalent to the Diablo member (Unit 5). Dibblee (1954) mapped Alverson volcanics and Split

Mountain Formation west of the Elsinore Fault; these were not examined but are probably part of the syn-rift nonmarine rocks (Unit 2).

Jactmlba area and Volcanic Hills (Figure 7.2)

Near Jacumba, the stratigraphie sequence consists of the Table

Mountain Gravels, a pre-rift unit (1), and the Jacumba Volcanics, which includes Miocene tholeiitic to alkalic basalts (Minch and Abbott 1973,

Hawkins 1970), and is apparently a syn-rift unit (2). Fourt (1979) described the following stratigraphie succession in the Volcanic 335 Hills: (1) a thin basal conglomerate, the "Split Mountain Formation", possibly equivalent to the Red Rock member in the Coyote Mountains

(Unit 1); (2) olivine basalt of the "Alverson Canyon Formation" (Unit

2); and (3) fine sand and mud of the "Palm Springs [sic] Formation"

(Unit 5) and Canebrake conglomerate (Unit 7). This sequence is similar to that observed by the author in the western Coyote Mountains east of Canyon Sin Nombre.

Sierra de los Cucapas (Figure 7.3)

The only detailed description of Neogene stratigraphy in this area is that of Barnard (1968), who described most of the same genetic units that occur elsewhere in the western Salton Trough. Apparently, no pre-rift clastics (Unit 1) occur in this area. Barnard's "older fanglomerates" and Colonia Progreso volcanics (dated at 15.3 + 0.8 Ma) represent the syn-rift phase (Unit 2). The reported zone of interfingering between his "older fanglomerates" and Imperial Formation is similar to the Latrania member, and represents the pre-deltaic marine phase (Unit 3). The bulk of his Imperial is similar to the Mud Hills and DeGuynos members, and constitutes the deltaic marine sequence (Unit

4). His Palm Spring Formation is similar to the Diablo member and represents the Colorado delta plain (Unit 5). Basin-margin alluvium

(Unit 7) is represented by his Canebrake Formation.

This area has special importance in that these are the southernmost known occurrences of the Neogene deltaic sequence, when strike-slip displacements are corrected for (Chapter 8). To the south, the nearest Neogene marine localities (near San Felipe) show no 336 evidence of Colorado River-derived sediments or deltaic facies (Boehm 1982, 1984), although Holocene Colorado-derived beach sand extends nearly to San Felipe (Kniffen 1932). The delta-plain facies in the Sierra de los Cucapas is somewhat atypical in that its color is report- more edly drab than in the typical Diablo member, and gravel derived from the Colorado Plateau is present (Barnard 1968). In the typical

Diablo member, Colorado River-derived sediment is no coarser than medium sand (Chapter 5), and the modern Colorado River downstream from Yuma carries only sand, silt, and clay (Barnard 1968). The Cucapas

locality is clearly more distal with respect to the ancestral Colorado

delta plain than is the FCV section, so the gravel-bearing delta-plain sequence in the Cucapas must be younger than the Diablo member in the

FCV section (4.0 to 2.8 Ma). The anomalous gravel in the delta-plain sequence of the Cucapas may indicate a hydrologic change in the Color-

ado River during the late Pliocene or Pleistocene, perhaps related to glaciation in the headwaters of the Colorado River (Barnard 1968).

Cerro Prieto Field

Of several Salton Trough geothermal fields located on the

Colorado delta plain aucchitta 1979,) Cerro Prieto is the only one with extensive published data. Ingle (1982) examined microfossils from

14 core samples from 185 to 1952 m below the surface. Foraminifers were found in a shaly unit between 700 and 1100 m, which he assigned to the mid-Pleistocene. In most other samples ostracodes were found but no in situ foraminifers, suggesting fresh or brackish-water deposition. Ingle

(1982) tentatively placed the Plio-Pleistocene boundary near a depth of 337 2000 m, based on the occurrence of Elphidium gunteri, which he regarded as a Pliocene indicator. In summary, the Cerro Prieto section appears to represent the Plio-Pleistocene delta plain (Unit 5) with a marine tongue (Unit 4).

Superstition Mountain

Dibblee (1954, 1984) mapped a local outcrop of Split Mountain

Formation, Alverson volcanics, and Imperial Formation in Superstition

Mountain. On field reconaissance of this area, the author was unable to

find the Imperial; however a small exposure of conglomerate, basalt,

and tuff was observed, confirming the section described by Ruisaard

(1979). Gjerde (1982) reported a date of 16.1 Ma on alkali basalt from

this locality, supporting a correlation with the Alverson volcanics in

the Coyote Mountains. All of the Neogene rocks observed at this local-

ity are assigned to the syn-rift nonmarine phase (Unit 2).

Whitewater Canyon and vicinity (Figure 7.4)

No evidence of the deltaic sequence (Units 4 and 5) was

observed in this area. The Coachella Conglomerate is almost certainly a

syn-tectonic fanglomerate (Unit 2); Peterson (1975) reported a date of

10.0 + 1.2 Ma obtained by D. Krummenacher from a pyroxene andesite autobreccia in the lower part of the Coachella. A thin unit of Imperial

Formation similar in appearance to the Latrania member overlies the

Coachella with angular discordance; sandstones within the Imperial are entirely of the "L" suite and the siltstone facies is very micaceous,

SO it is regarded as a pre-deltaic marine unit (3). The overlying 338 Painted Hill Formation consists of conglomerates and coarse sandstones and is considered a basin-margin unit (7). Matti et al. (1985) reported K-Ar dates of ca. 6.0 Ma from a basalt flow a short distance above the top the the Imperial. The Neogene sequence north of Banning was not examined in the field, but on the basis of published descriptions

(Bramkamp 1935, Proctor 1968) it apparently consists of the same unit as the Whitewater locality.

Indio Hills (Figure 7.5)

An equivalent to the Mud Hills member (Subunit 4b) is exposed near Edom Hill, and a Deguynos equivalent (Subunit 4c) is exposed at

Willis Palms. A unit of marine mudstone reported from the Texas 1-Edom well and assigned to the Imperial Formation is probably also part of the deltaic sequence (Unit 4). All the other strata observed in the

Indio Hills are "L"-suite sandstones and conglomerates, and are assigned to the nonmarine basin-margin unit (7). Stotts (1965) reported an angular unconformity within the nonmarine clastics, between his

Thousand Palms and Indio Hills Formations. His report of a Pleistocene unit below the marine mudstones is discounted as unlikely, in view of the structural complexity of this area.

Durmid Hill (Figure 7.6) and vicinity

Babcock (1969, 1974) reported two units from this area: a nonmarine coarse-clastic unit (7), which he named the Shavers Well

Formation, and a lacustrine unit (6) which he assigned to the Borrego

Formation. Merriam and Bischoff (1975) interpreted an ash bed within 339 the lacustrine unit as an equivalent of the 0.7 Ma Bishop Tuff of east-central California. However, Sarna-Wojcicki (personal communication to Wagoner 1977) proposed a correlation with the 0.6 Ma Friant

Pumice in the San Joaquin Valley.

Fortuna and San Luis Basins (Figure 7.7)

Neogene strata in the Yuma Desert are not exposed; stratigraphy of this area is known only from well logs and cuttings. The lowest sedimentary sequence encountered in wells consists of nonmarine coarse clastics (Olmsted et al. 1973; Eberly and Stanley 1978), presumably syn-rift fanglomerates (Unit 2). The unit is unconformably overlain by marine sandstone and mudstone; as discussed in Chapter 4, this is probably equivalent to the Wind Caves member (Subunit 4a). It is not

apparent whether an equivalent to the Latrania or Lycium members occurs

in any of these wells, although the basal marine limestone reported in

wells LCRP-26 and LCRP-23 (Olmsted et al. 1973) could be construed as a

pre-deltaic marine unit (3). The Bouse Formation reported in these

wells is apparently equivalent to the Mud Hills member (Subunit 4b) and

possibly to the Deguynos member (Subunit 4c); the overlying transition-

al unit is placed in Subunit 4d. "Older Colorado alluvium" (Unit 5)

overlies the "transitional unit". Locally-derived conglomerate and

sandstone (Unit 7) locally interfinger with the "older alluvium".

Bouse EMbayment (Figure 7.8)

The Bouse Formation is underlain by nonmarine coarse clastics are (Metzger et al. 1973, Metzger and Loeltz 1973, Buising 1986); these 340 probably related to Basin-Range faulting and could be construed as a syn-rift unit (2), although not actually part of the proto-Gulf rift basin. In the Bouse Embayment, the marine-to-brackish Bouse Formation consists of a basal tuffaceous marine unit with some locally-derived clastics (Unit 3), and an overlying "interbedded unit" (Metzger et al.

1973, Metzger and Loeltz 1973). Sands in the interbedded unit, observed

in cuttings of the Bouse type well (LCRP-27) are predominantly of the

suite; therefore, this unit is considered as a deltaic marine

sequence (Unit 4), but a direct correlation with facies of the FCV

section was not possible. Terraces (Unit 8) and the modern floodplain

of the lower Colorado River are entrenched below the highest occurrence

of the Bouse, indicating an unconformity separating the marine and nonmarine Colorado-derived sediments.

Possible extensions of Bouse Embayment

A number of deposits of clay, limestone, and evaporites

peripheral to the Bouse Embayment have been suggested as possible

marine, estuarine (brackish-water), or hypersaline extensions of the

Gulf of California. Of the following, only the occurrences at Danby and

Cadiz Dry Lakes have strong fossil evidence for a marine connection; others may represent saline nonmarine environments. Possible extensions

of the Bouse Embayment in Arizona have most recently been reviewed by

Scarborough (1984).

Mojave Desert. An extension of the Bouse Embayment west into the Mojave was suggested by core descriptions from test wells at Danby and Cadiz Dry Lakes (Bassett, Kupfer, and Barstow 1959) and confirmed 341 by the reported presence of Globigerina in a core from the Danby-1 well (Durham and Allison 1960). Microfaunas from these wells have also been discussed by Smith (1960, 1970), Metzger (1968), and Winterer (1975). The Danby wells yielded six species of foraminifers, while the

Cadiz well yielded only Ammonia beccarii; cores from both dry lakes contained fossils of barnacles and molluscs (Smith 1970). Cores from

Bristol Dry Lake showed similar lithologies to those of Danby and

Cadiz, including evaporites, but no marine microfossils (Bassett et al. 1959)

Microfossils were also found in the Panamint Core-1 well at

Panamint Dry Lake, including Elphidium, ostracodes, diatoms, and Chara sp. (Smith 1960, Smith and Pratt 1957, Winterer 1975). The presence of

Elphidium suggests salinities of 32-37 o/oo (near normal marine salinities), but these microorganisms could have been introduced into a saline lake on the feet of birds, so a marine connection is not required (Winterer 1975).

Lake Mead area. The predominantly lacustrine Muddy Creek Forma- tion (Longwell 1936) has long been recognized to predate initiation of the Colorado River in the Lake Mead area (Longwell 1946); in this regard the Muddy Creek is analogous to the Bouse Formation. The

Bualapai Limestone, which in places attains a thickness >300 m, is essentially the upper member of the Muddy Creek (Blair 1978). Blair

(1978) proposed a marine or estuarine origin for the Hualapai on the basis of cristobalite lepispheres found in chert, and carbon-isotope data from limestone. Cornell (1979) argued that either brackish or 342 hypersaline conditions could explain these observations, and that a marine connection is not required. Blair and Bradbury (1979) stated that microfauna and microflora in the Hualapai rule out hypersaline environments. Lucchitta (1979) favored deposition in a saline terrestrial lake without a marine connection, noting a large gap between the geographic distributions of the Bouse Formation and Hualapai Lime- stone. He also noted that the Muddy Creek Formation has an evaporitic facies, indicating hypersaline conditions.

Age of the Hualapai Limestone is constrained by three radiometric ages: 10.9 + 1.1 Ma on a lava flow at the base of the Hualapai

(Blair 1978); 8.7 + 2.2 Ma on an air-fall tuff within the Hualapai

(Blair 1978); and 5.8 + 0.2 Ma on the Fortification Basalt, which directly overlies the Hualapai (Shafiqullah et al. 1980). On the basis of these dates, it appears that the Hualapai predates the Bouse transgression, thus lending support to a nonmarine origin as preferred by

Cornell (1979) and Lucchitta (1979).

Gila Valley. Ross (1923), who originally described the beds that would later be called Bouse (Metzger 1968), also noted the presence of subsurface clays along the Gila River extending upstream to

Gila Bend, and suggested a lacustrine or estuarine origin. Metzger

(1968) suggested that these clays represent an extension of the Bouse

Enamtyment,' although no marine microfossils have been reported.

Upstream from Gila Bend, in the Phoenix area, is an extensive topographic low ("Gila Low" of Peirce 1976) underlain locally by thick evaporite deposits (Eaton, Peterson, and Schumann 1972, Peirce 1976, 343 Eberly and Stanley 1978). The Picacho anhydrite body, which may be as much as 2 km thick, overlies a volcanic unit dated at 15 Ma (Peirce 1976, Shafiqullah et al. 1980). The Luke halite body, possibly 4 km

thick, includes a basalt flow near the top which has been dated as 10.5 Ma (Shafiqullah et al. 1980). Therefore, the Luke and possibly the

Picacho apparently pre-date the Bouse Transgression, which lends sup-

port to a nonmarine origin. Initiation of the lower Gila River is

bracketed by an 8.9 + 0.3 Ma basalt flow which underlies Gila River

gravels near Florence (Shafiqullah et al. 1980), and a 6 Ma basalt flow

which overlies Gila River gravels near Gillespie Dam in the Gila Bend

area (Scarborough 1984). If these dates are correct they indicate that

the Gila River was initiated before or during the Bouse transgression,

which might rule out a Bouse extension in the Gila valley.

Circum-Gulf localities

A number of important Neogene marine localities of the

circum-Gulf were apparently beyond the direct sedimentary influence of

the Colorado delta, so that the system of genetic-stratigraphic units

used for the Salton Trough is inappropriate. These localities are

nevertheless important for establishing the transgressive history, paleobathymetry, and oceanography of the early Gulf.

Sierra San Felipe and Santa Rosa

A Neogene marine sequence near San Felipe in northern Baja

California has recently been studied by Boehm (1982, 1984). The base of the sequence is not exposed, but the lowest unit consists of laminated 344 diatomite with a bathyal microfauna (San Felipe Diatomite Member).

Boehm (1984) gave a biostratigraphic age of 5.5-6.0 Ma (latest Miocene)

for this unit. The diatomite indicates the presence of a deep-marine

basin with a well-developed oxygen minimum and vigorous upwelling by

the end of the Miocene (Boehm 1984, Ingle 1984). The diatomite grades

up into more clastic-rich, shallower-water and more aerobic deposits

(Canon las Cuevitas member). That unit is overlain with angular uncon-

formity by the Pliocene marine coarse clastics of the Salada Formation

(middle to late Pliocene).

Isla Tiburem

A rich shallow-water calcitic macrofauna similar to that of the

Latrania member has been described from a thick (>2 km) sequence of

marine and nonmarine clastics which grades up into andesitic volcani-

clastics (Smith et al. 1984, Stump 1981). The following radiometric

ages have been reported: (1) 14.0 + 3.0 Ma on a volcanic flow within

the marine unit (Gastil 1975, cited in Stump 1981); (2) 12.9 + 0.4 Ma

on a volcanic clast within a debris flow (Smith et al. 1984); and (3)

11.2 + 1.3 Ma on an overlying ash-flow tuff (Smith et al. 1984). is Therefore, this unit is definitely middle Miocene. The macrofauna

dominated by pectinids, giant Ostrea, Turritella, barnacles, and

echinoids (Smith et al. 1984).

"Hermosillo Basin" (new name)

Gómez (1971) listed a planktonic foraminiferal fauna from

cuttings from water wells drilled near Hermosillo, Sonora. He 345 interpreted a middle Miocene and possibly early Miocene age for these faunas. However, the species listed by G niez are not diagnostic of the reported ages, and samples have been unavailable so that identifications could not be confirmed (M. C. Boehm 1986, oral communication).

Coastal Baja California

Durham (1950) listed a number of shallow-marine Pliocene local-

ities from the east coast of Baja and adjacent offshore islands, from

Isla Monserrate to Isla Angel de la Guarda. He referred to these

collectively as the Salada Group, and subdivided them into lower,

middle, and upper Pliocene units on the basis of their macrofaunas.

These faunas are dominated by pectinids, Ostrea, Encope, and

Clypeaster, similar to the Latrania fauna.

San José del Cabo Trough

McCloy (1984) described a Neogene marine sequence from this

locality near the southern tip of the Baja peninsula. The base of the

marine sequence consists of middle to upper Miocene inner-neritic

sandstone, siltstone, and mudstone; these grade up into finer-grained,

deeper-water sediments, culminating in Pliocene bathyal mudstones and

diatomite and siltstones. The oldest marine strata at this locality may

be older than those around the modern Gulf mouth, described below. If

SO, the San José del Cabo Trough may represent the original mouth of

the Gulf. 346 Cabo Corrientes

Neogene marine strata have been reported from a few localities in and around the Gulf mouth. Jensky (1974) described stratigraphy of

the Cabo Corrientes area on the Mexican mainland, on the southern side

of the Gulf mouth. Most of the Neogene rocks there consist of nonmarine

volcanics, conglomerates, and sandstones, but marine mudstone and

claystone with sandstone and conglomerate occur at Punta Mita. These marine beds reportedly contain an upper Miocene (Delmontian and/or

Mohnian) benthic foraminiferal fauna. The marine strata overlie and

interfinger with basalt flows and pillow basalt; a K-Ar date of 10.2 +

0.8 Ma was obtained from a flow. Jensky (1974) also reported K-Ar ages

from other mafic rocks in the area: 8.3 + 0.6 Ma on basalt, 10.1 + 0.3

and 13.3 + 2.0 Ma on gabbros, and 12.6 + 0.4 Ma on a mafic dike.

Tres Marias Islands

Geology of the Tres Marias Islands in the Gulf mouth is poorly

known. Jensky (1974), citing Chinas (1963), mentioned upper Miocene to

Pleistocene marine sediments on Maria Magdalena Island. Ingle (1984)

reported that upper Miocene diatomites occur on Marla Madre Island.

Mouth of Gulf

Neogene marine sediments were also recovered from DSDP sites

474, 475, and 476, on the northern side of the Gulf mouth (Currey et al. 1982). At site 474, located on oceanic crust, a sequence of muddy turbidites with diatoms and nonnofossils overlies pillow basalts.

Nennofossils indicate an age at least as old as Zone NN16, and possibly 347 NN16 (lower Pliocene); magnetic anomalies suggest a crustal age of 3.2 Ma. At sites 475 and 476, on block-faulted and attenuated continental crust, similar stratigraphic sequences were reported, with nannofossils indicating an age no greater than 3.7 Ma. At site 476, however, two diatom species suggested an age of 4.4 to 4.5 Ma. A fresh basalt cobble from Hole 475 B yielded a K-Ar age of 10.2 + 2.1 Ma, similar to ages reported by Jensky (1974) from the matching southern side of the Gulf mouth.

Regional chronology

Salton Trough and Bouse Embayment

Several radiometric ages constrain the timing of marine transgression and regression in the Salton Trough and Bouse Embayment

(Figure 7.1). In the Coyote Mountains and Carrizo Impact Area, the

Alverson volcanics dated at 14-22 Ma underlie the lowest marine strata of the Imperial Formation (Latrania member), which are typically separated from the Alverson by an angular unconformity, nonmarine interval, or both. Therefore, marine transgression may be significantly later than 14 Ma. Because marine inundation was almost certainly diachronous, these dates do not necessarily constrain the age of the oldest marine strata at Split Mountain, which are not underlain by the

Alverson. For the FCV section, the timing of marine regression by deltaic progradation is tightly constrained at 4.0 Ma by magnetostratigraphy (Figure 6.3).

Marine transgression and regression near Whitewater Canyon

(Super Creek) are constrained by dated basalts in nonmarine units that 348 bracket the Imperial Formation: 10 Ma in the underlying Coachella Conglomerate (Peterson 1975) and 6.0 Ma in the overlying Painted Canyon Formation (Matti et al. 1985). The 10 Ma basalt is separated from the Imperial Formation by an angular unconformity and by a thicker nonmarine section than is the 6.0 Ma basalt. This suggests that the age of the Imperial is probably closer to 6 Ma than 10 Ma.

In the Bouse Embayment (Figure 1.2), dates of 5.5 Ma from tuff

in the base of the Bouse Formation at Milpitas Wash (Damon et al. 1978,

Shafiqullah et al. 1980) place a minimum age on marine transgression.

However, this locality probably represents only the fringe of the

original Bouse basin; the thickest Bouse and the deepest part of the

basin occur under the Colorado floodplain (Metzger et al. 1973, Metzger

and Loeltz 1973, A. K. Buising 1986, oral communication). Therefore,

the axis of the basin was probably inundated earlier than the Milpitas

Wash locality. A date of 3.8 Ma on a basalt overlying a low terrace

deposit of the Colorado River (Damon et al. 1978, Shafiqullah et

al. 1980) places a minimum age on marine regression from the Bouse Embayment.

These dates of 3.8 and 5.5 Ma also bracket uplift of the Bouse

Embayment (Lucchitta 1972, 1979) and initiation of the lower Colorado

River. Consequently, the 5.5 Ma date also provides a maximum age on the first influx of Colorado River sediment in the Salton Trough, and therefore on the base of "C"-suite units ("deltaic sequence") in the

Salton Trough (Figure 7.1). This result is consistent with ages from the FCV section discussed in Chapter 6; the base of the lowest 349 "C"-suite unit (Wind Caves member) is >4.3 Ma by magnetostratigraphy, and this unit was assigned to foraminiferal zone N19 (<5.1 Ma) by Stump (1972).

The magnetostratigraphic age of 4.0 Ma on the Imperial-palm Spring contact in the FCV section marks the first time at which the nonmarine Colorado delta plain definitely extended across the Salton

Trough to the western margin. Therefore, all marine units farther north in the Salton Trough must be older than 4.0 Ma (except in the unlikely event that the northern Salton Trough had a marine connection with the

Los Angeles Basin). This result is consistent with dates from Whitewater Canyon.

Paleocurrent data (Chapters 4 and 5) and paleogeographic reconstructions (Chapter 8) clearly indicate that the ancestral

Colorado delta on the western margin of the Salton Trough prograded from north to south. Therefore, the date of 4.0 Ma on the Palm Spring-

Imperial contact in the FCV section is also a maximum age on the contact between Units 4 and 5 in the Yuha Basin, and a minimum age on the same contact in the San Felipe Hills (Figure 7.1). Barnard's (1968) report of Colorado Plateau-derived gravel in the delta-plain sequence of the Cucapas suggests that this unit is younger than the delta-plain sequence of the FCV section (4.0-2.8 Ma), which contains no such gravel; therefore, Unit 5 in the Cucapas is tentatively dated as <2.8 Ma.

The age of lacustrine units is generally not well constrained. Babcock (1974) estimated ages of 0.6 to 0.7 Ma for an ash 350 bed in the "Borrego Formation" at Durmid Hill. However, this lacustrine unit is not necessarily coeval with the type Borrego in the Borrego Badlands or San Felipe Hills, which is generally regarded as Pliocene

(Dibblee 1984). Work in progress on paleomagnetism (Bogen and Seeber

1986) and vertebrate stratigraphy (George Miller 1985, oral communica-

tion) of the Borrego Formation and overlying Ocotillo Conglomerate may

help to constrain the age of the Borrego.

These data provide an internally consistent Neogene chronology

of the Salton Trough and Bouse Embayment, including marine trans-

gression, uplift of the Bouse Embayment, initiation and incision of the

lower Colorado River, initiation of the Colorado delta in the Salton

Trough, and isolation of the Salton Trough from the Gulf of California.

These events occurred within a short period of time, perhaps within few

million years, probably too rapid to be subdivided by biostratigraphy.

Marine units in the Salton Trough and Bouse Embayment straddle the

Miocene-Pliocene boundary at 5.1 Ma, but that boundary cannot be

located precisely within individual stratigraphic sections.

Circus-Gulf of California

Volcanic rocks of the circus-Gulf region place important

constraints on the tectonic history of the early Gulf (Karig and Jensky

1972). Radiometric ages for volcanic rocks and other datable phenomena were plotted for the last 20 Ma, in geographic order from north to south (Figure 7.9). In Baja California Norte and Sonora, peak Miocene volcanism coincided with an episode of east-west extension and drainage reversal from west to east (Dokka and Merriam 1982, Gastil and

351

m < cc ct,,c0 Lii ...LI, I. i•-• z cr; cc • 40-52 0. 02cc z co

Cf; 0 2 I- (7) co CO Z -I 2 w 2 LU < ?_ cocci U_ 0 0 Z '- 'Z — 0 -1 LIPO 0 z 0 co cc To = D W al > WCOZ Z.t7 .-....-, Z 17-.c U.1 Lij < 11 < 7; XI-< Cc Q-J l'rr) < ..., t u j co — Z M (i) ci) 0) CC '', 0 0 w z u) I _1- > (5 1- z w 0 w • co Lu co

u_ Fc

u_ Fc O

n n,,,....• ...v../ ••n•••.,- ..../..... •••n...... -/ •••n•••„.n•••••• •... . N N .4. CO CO O-;' -Y co 73 GO CO o as c) c0 0 -c ..5 0.) cp co .ct - o) c.) ,-- c .- - cul•••- co ii; 0) co ''' .o co CD 7-; c N. N. co a, co coo, :-.: a, 0 _ 8 >, c- -“ ï 2 0 co co m E — 0 co co — co — co = c0 z 2 t.- 0E O ' CO .E c < ...c 8 (..1 Cl) 0 Y 2 352 Krummenacher 1977). Similarly, in the FCV area the onset of Miocene alkalic volcanism was roughly coincident with initiation of fanglomerate deposition, normal faulting, and drainage reversal (Chapter Baja California 6). In Sur, a well-defined change from intermediate, calc-alkaline, subduction-related volcanism, to mafic, alkalic to

tholeiitic, rift-related volcanism has been documented by Sawlan and

Smith (1984) and Hausback (1984). Their dates are consistent with those

of Jensky (1974) for mafic igneous rocksnear Cabo Corrientes, pre-

viously discussed. In the Colima graben, southeast of the Gulf mouth,

alkalic volcanism began at -5 Ma, considerably later than in Baja California or Cabo Corrientes.

Volcanic dates from the Chocolate Mountains and Mojave block

are not as directly related to circum-Gulf chronology and are not plotted in Figure 7.9, but are interesting to compare with ages of the

Alverson basalts. Crowe, Crowell, and Krummenacher (1979) reported a termination of subduction-related, silicic to intermediate volcanism in the Chocolate Mountains at -22 Ma, about the same time as the earliest

Alverson basalts were formed. In the Mojave block, farther to the north, a major change from predominantly silicic volcanicm to predominantly mafic, alkaline volcanism occurred from 18 to 15 Ma (data compiled by Gary Calderone 1986, oral communication).

In summary, a change from subduction-related to rift-related volcanism occurred first in the Chocolate Mountains and western Salton

Trough, at -22 Ma, and occurred somewhat later to the north and south. A crude north-to-south diachroneity is observed from the Salton 353

Trough to the Colima Graben (Figure 7.9), but the onset of rifting was approximately synchronous in much of the circum-Gulf region, at -14-12 Ma.

Marine inundation probably proceeded from south to north, based on tropical Caribbean faunal affinities in the Gulf of California

(Durham, 1954). Marine waters reached Isla Tiburcin by 12 Ma (Smith et al. 1985), Sierra San Felipe by 6 Ma (Boehm 1984), Whitewater Canyon before 6 Ma (Matti et al. 1985), and the Bouse Embayment by 5.5 Ma

(Shafiqullah et al. 1980). CHAPTER 8

TECTONIC DISPLACEMENTS, PALINSPASTIC RECONSTRUCTION, AND PALEOGEOGRAPHY

The greatest problem in paleogeographic reconstruction of the Salton Trough region is to develop a satisfactory series of palinspastic base maps to account for regional syn-depositional deformation. To develop such base maps, two approaches were used. The first approach was to consider a geometrically simple reference model that satisfied most of the published constraints, and to plot resulting tectonic trajectories (Winker and Kidwell 1986). The second approach was an attempt to quantify the uncertainty on reconstructed positions of individual localities, so that these positions could be plotted with appropriate error bars and the paleogeographic implications of these uncertainties could be considered.

Within the Salton Trough region direct constraints on tectonic displacements are few, so it is necessary to incorporate data from outside the immediate study area. This analysis is based on the following assumptions; (1) regional deformation can be treated to a first approximation as relative translations of seven rigid blocks (Figure

1.1); (2) the history of the San Andreas fault system is closely related to northwest-southeast opening at the mouth of the Gulf of

California; and (3) all significant faults have been identified. The

354 355 possibility of internal deformation of blocks, such as by simple shear, extension, contraction, or block rotation, can be considered later as deviations from the rigid-block approximation.

Reference model

Moore and Curray (1982) proposed a model for the motion of Baja

California relative to the North American Plate (NAP) based in part on sea-floor magnetic anomalies at the mouth of the Gulf of California

(Plate 1). They postulated that Baja California translated to the northwest relative to the NAP at a constant rate of 56 mm/a (Minster and Jordan 1978) beginning at 5.5 Ma, for a total displacement of -310 km. The early displacement is thought to have been accommodated entirely by stretching of continental crust (rift phase). The beginning of sea-floor spreading (drift phase) was diachronous, starting possibly as early as 4.7 Ma and definitely by 3.9 Ma, and involving the entire Gulf mouth by 3.2 Ma. Displacement history is directly constrained by magnetic anomalies only for the drift phase.

Displacement between Baja California and the NAP was accommodated in the northern Gulf-Salton Trough by a number of strike-slip faults in a broad zone from the San Andreas fault (sensu strictu) to the Agua Blanca fault (Figure 1.1, Plates 1, 2). Most of this displacement was probably accommodated by a relatively small number of major faults (Crowell and Ramirez 1979). Therefore, displacement histories are treated in terms of seven rigid blocks (Figure 1.1). Straight,

Parallel trajectories are also assumed, and an azimuth of 315° is used, based on the dominant trend of the major faults. Since displacement 356 histories for individual faults are not known in detail, constant rates of displacement are assumed in the reference model, each starting at

5.5 Ma and therefore proportional to the total displacement. The following total displacements (rounded to the nearest 5 km) are used: (1) 20 km for the Agua Blanca fault (Allen, Silver, and Stehli 1960); (2) 40 km for the Elsinore fault (Sage 1973, Crowell and Ramirez 1979); (3) 25 km for the San Jacinto fault (Sharp 1967); (4)

215 km for the Mission Creek branch of the San Andreas fault (Peterson

1975). This yields a total displacement of 300 km; the discrepancy between this total and the required 310 km is presumed to be accommodated by the Banning brach of the San Andreas fault, for which the total displacement has not been determined (Allen 1957 suggested a minimum of 7 km). In the present kinematic system, strike-slip displacement on the San Andreas fault in the Salton Trough region is offset by a series of crustal spreading centers to the Imperial fault and then to the Cerro Prieto (San Jacinto of Merriam 1965) fault

(Crowell and Ramirez 1979, Rockwell and Sylvester 1979, Elders 1979b,

Fuis and Kohler 1984). Long-term histories of the Imperial and Cerro

Prieto faults are unknown, but they are not critical for Neogene reconstruction.

The Algodones fault system (Mattick et al. 1973, Olmsted et al.

1973) has been suggested as a former southeastern extension of the San

Andreas fault (Crowell and Ramirez 1979), although strike-slip displacement has not been demonstrated. If a displacement of 215 km is accepted for the Mission Creek fault, an overlap between the Indio 357 Hills and San Luis Basin localities would result unless some strike- slip displacement on the Algodones fault is included as well. With these assumptions, tectonic trajectories can be plotted for all key Neogene localities except for the San Luis Basin, which is unconstrained by field evidence. The palinspastic base can then be used to construct a series of paleogeographic maps (Figure 8.1). These

reconstructions will be discussed after consideration of some of the

inherent uncertainties in constraints on the palinspastic model.

Constraints on tectonic displacements

Pacific-North American plate motion

Minster et al. (1974) calculated a globally consistent model of

instantaneous relative plate motions using a linearized maximum-

likelihood inversion of sea-floor spreading rates, orientations of fracture zones, and earthquake slip vectors. Spreading-rate data from the mouth of the Gulf of California were specifically excluded because of the possibility of a separate Rivera Plate (Plate 1) that is not coupled to the North American Plate (Atwater 1970, Molnar 1973). The orientation of the San Andreas fault was also excluded from the data set because extension in the Basin-and-Range Province may also contri- bute to the relative plate motion. Minster et al. (1974) obtained a slip rate of 55 mm/a between the North American and Pacific Plates.

Minster and Jordan (1978) recalculated the model with an improved data set, and obtained a revised slip rate of 56 + 3 mm/a.

Atwater and Molnar (1973) calculated a series of late Cenozoic finite slip rates between the North American and Pacific Plates, using •

358

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•P • 1., O (0 • CO Cr) G — u° .4 —I to o -a 1.4

O g .4) au oo vz a) c+••n O .e•I

0 0 • z+r $.n • 4.) 0) 4% 0 O U (.) c.) c .= a eu 4-1 s. s•-, o a) roc o

cl a, a, a

G) •,1 0) tO a, o a) c z 0

-.I a) cc; 2 Q) 5. ai

• W.n 0

-4 r••nI

C.) 359 sea-floor magnetic anomalies to constrain a vector circuit of the North American, African, Indian, Antarctic, and Pacific Plates. Blake et al. (1978) published corresponding azimuths for the same periods of time discussed by Atwater and Molnar; combined results of these two studies are tabulated in Appendix IV and depicted as vectors in Plate 1. For example, at 330 latitude, near the Salton Trough, the calculated finite motions are 13 mm/a at 3350 from 21 Ma to 10 Ma, 45 mm/a at 326°

from 10 Ma to 4.5 Ma, and 55 mm/a at 319° from 4.5 Ma to the present.It

should not be inferred, however, that instantaneous motions were neces-

sarily constant during these periods, or that plate motions changed abruptly at 10 Ma and 4.5 Ma.

Spreading at mouth of Gulf

Maps of sea-floor magnetic anomalies in and adjacent to the Gulf mouth have been published by Larson (1972), Mammerickx (1980),

Ness et al. (1981), and Curray et al. (1982), each using a different

data set; the last two of these studies are inc.:,,--porated in Plate 1.

Interpretation of magnetic anomalies back to -3.5 Ma is fairly

straightforward, but interpretation of older anomalies has been con-

troversial. The Rivera Ridge spreading center is located asymmetrically

Within the Gulf mouth, and older, pre-3.5 Ma oceanic crust occurs west

Of the Islas Tres Marias. Mammerickx (1980, and Klitgord 1982) attrib-

uted the older crust to pre-3.5 Ma opening of the Gulf around a now- extinct spreading center, followed by a ridge jump at 3.5 Ma. In contrast, Ness et al. (1981) interpreted the older crust in terms of

Spreading about the present Rivera Ridge. They attributed the 360 asymmetric location of the Rivera Ridge to northwesterly subduction of older oceanic crust under the tip of Baja California. Moore and Curray (1982) rejected both of these interpretations. They regarded magnetic mummdies immediately west of the Islas Tres Marias as ambiguous and insufficient to constrain the age of the older oceanic crust. In their model, crustal extension at the Gulf mouth began -5.5 Ma, and oceanic crust began to form as early as 4.7 Ma; by 3.2 Ma the Gulf mouth was opening purely by sea-floor spreading. Regardless of the interpretation of the older oceanic crust, spreading rates for the last -4 Ma inferred from the magnetic anomalies (Figure 8.2) are essentially indistinguishable from the rate of 56 + 3 mm/a between the North American and

Pacific Plates obtained independently by Minster and Jordan (1978).

These anomalies do not necessarily constrain the motion between

North America and Baja California, however, because of the possibility of a Rivera Plate distinct from the North American Plate (Atwater

1970). Molnar (1973) showed that well-determined focal mechanisms for the Rivera Fracture Zone indicate a rotation pole significantly different from the North American-Pacific pole, which would require a separate Rivera Plate. Ness and Lynn (1975) pointed out that the curvilinear trend of the Rivera Fracture Zone is incompatible with either the North American-Pacific or Cocos-Pacific Euler poles of relative rotation. They interpreted this trend to result from ongoing transfer of the Rivera Plate from the Cocos Plate to the North American

Plate, and suggested that "initial opening of the mouth of the Gulf of 361

180

north of Rivera Ridge

160 — • (Curray et a1,1982)

north of Rivera Ridge (Ness et a1,1981) 140

tr) n south of Rivera Ridge

1-1 (Ness et a1,1981)

CD 8. 100 Cl)

Et" E 60 o

o co 40

\\ Pacific-North American 20 \ plate motion (independently derived) (Minster & Jordan, 1978)

1 2 3 4 5 6 Age (Ma)

Figure 8.2. Age vs. distance for magnetic anomalies in mouth of Gulf of California. 362 California may have occurred along lines oblique to North American- Pacific relative motion."

The compatibility of spreading rates measured directly from the Gulf mouth with Pacific-North American slip rates calculated indirectly from an independent global data set suggests that the Rivera Plate is

essentially coupled to the North American Plate and Baja California is

essentially coupled to the Pacific Plate, at least within a few mm/a.

This two-plate approximation will be assumed in developing the palins-

pastic model. However, it is geometrically possible that significant

right slip occurs between Baja California and the Pacific Plates and an

equivalent amount of right slip occurs between the Rivera and North American Plates.

If the two-plate approximation is valid, magnetic anomalies in

the Gulf mouth provide a tight constraint on the motion of Baja Cali-

fornia relative to the North American Plate from -4 Ma to the present.

Moore and Curray (1982) extrapolated the essentially constant spreading

rate to 5.5 Ma to account for the -300 km displacement on the San

Andreas fault. However, the total offset at the Gulf mouth is not well

determined and may be significantly greater than 300 km, and the timing

of initial northwest-southeast extension is not known. In fact, radiometric ages from mafic igneous rocks around the Gulf mouth (Chapter 7)

suggest that crustal extension may have occurred as early as 13 to 8 Ma (Jensky 1974).

Two lines of evidence suggest that right slip between the

Pacific and North American Plates was initially accommodated outboard 363 of Baja California. Spencer and Normark (1979) identified the Tosco- Abreojos fault zone on seismic-reflection profiles offshore from southern Baja California (Plate 1). Yeats and Haq (1981) demonstrated that the Magdalena Fan west of Baja California (Plate 1), which was deposited from 14.5 Ma to 13 Ma (+ 0.5 Ma), has apparently been dis-

placed 600-700 km since that time from its probable source in the nascent Gulf mouth, or substantially farther than the displacement of

Baja California. The extra slip was probably accommodated on the Tosco- Abreojos fault system.

San Andreas and San Gabriel faults

To a first approximation, the San Andreas fault can be regarded

as the transform boundary between the Pacific and North American

Plates, and its history should therefore be related to the opening of

the Gulf mouth. On closer examination, however, the San Andreas fault does not constitute a discrete plate boundary: (1) The San Andreas system includes a number of splays from the main fault, including branches which currently accommodate a significant amount of the relative plate motion (e. g. San Jacinto fault) or have done so in the past

(c. g. San Gabriel fault); (2) during the late Quaternary the San

Andreas fault has accounted for only -35 mm/a of right slip (Matti et al. 1985, Sieh and Jahns, 1984, Weldon 1984, Weldon and Sieh 1985,

Weldon and Humphreys 1986), or about two-thirds of the relative plate motion; and (3) the Transverse Ranges have had a complicated tectonic history, possibly including Miocene east-west to northeast-southwest

(il) present coordinates) extension (Campbell and Yerkes 1976), major 364 clockwise rotations during the late Cenozoic (Luyendyk, Kamerling, and Terres 1980, Luyendyk et al. 1985, Hornafius 1985, Kamerling and

Luyendyk 1985, Terres and Luyendyk 1985, Luyendyk 1986), and Quaternary north-south contraction (Matti et al., 1985). Because of this complex

tectonic history, outboard blocks north and south of the Transverse

Ranges may have significantly different displacement histories (Hornafius 1986).

With respect to the Salton Trough, the most important evidence on the San Andreas fault history comes from the vicinity of the San

Gabriel Mountains. The possibility of large strike-slip displacements on the San Andreas and San Gabriel faults adjacent to the San Gabriel Mountains was first suggested by Crowell (1952, 1962), who recognized striking similarities among basement rock assemblages in the Orocopia

Mountains (inboard of the San Andreas), Soledad Basin (between the two faults), and Tejon Pass (outboard of the San Gabriel). To juxtapose these terranes by restoring strike-slip offsets requires -60 km offset on the San Gabriel fault and -240 km offset on the San Andreas fault, for a total of -300 km, a figure that is widely quoted in the southern

California literature. Crowell's interpretation was strengthened by subsequent studies of the basement terranes and of the middle Miocene

Mint Canyon and Caliente Formations (reviewed in Crowell 1975a). Ehlig,

Ehlert, and Crowe (1975) discovered distinctive clasts of rapikivi- textured quartz latite porphyry in the nonmarine Mint Canyon Formation

(between the faults) and Caliente Formation (outboard of the San

Gabriel), and located a probable source in the Chocolate Mountains 365 (inboard of the San Andreas); these formations and their source area

show the full offset on the San Gabriel and San Andreas faults, which indicates that these units pre-date faulting. The Mint Canyon Formation contains Barstovian and Clarendonian vertebrate faunas; therefore, the

age of the Barstovian-Clarendonian boundary (12 Ma) places a maximum age on the San Gabriel and San Andreas faults.

Further evidence for the timing of these faults comes from

stratigraphic relationships in the Ridge Basin, exposed in the San

Gabriel Mountains between the faults (Figure 8.3, Crowell 1975 b, 1982,

Crowell and Link 1982, Link 1982, 1983, Ramirez 1983). The Violin

Breccia is a syntectonic unit along the San Gabriel fault; its lowest

part interfingers with the marine Castaic Formation of Mohnian (Pacific

coast provincial marine stage) age, which in turn overlies the nonmarine Mint Canyon Formation. Ensley and Verosub (1982) used bio-

stratigraphy and magnetostratigraphy to estimate an age of 8-9 Ma for

the upper Castaic Formation; this places a minimum age on initiation of

the San Gabriel fault. Outboard of the San Gabriel fault, the marine

Model° Formation, also of Mohnian age, is offset 30-35 km from its source area, or less than the total offset on the San Gabriel (Ehlig et al 1975, Crowell 1975b); this also requires that the San Gabriel fault was active during the Mohnian. The San Gabriel fault is overlapped by the nonmarine Hungry Valley Formation, which is cut by the San Andreas fault; no sedimentary evidence for earlier movement on the San Andreas has been observed in the Ridge Basin. Ramirez (1983) demonstrated that the upper part of the Hungry Valley, which has been tentatively dated 366

A. SW EXPLANATION RIDGE BASIN G5I Breccla Conglomerate

Apple. Can' n yo'n Santiotone— M rn.b;r.'. • ..••• M Sandstone • n ..4. p n •• k=7-1 SlItstone

• Shale (Mudstone) Basement Rocks

THICKNESS

17,000 4500 4000

101,000 3000 SAN — n FRANCISOUITO 2000 FORMATION 5000 1000 Feet— Meters O 1 2 3 Km. I O 1 2 Mlles Horizontal Scale 8.

2 I: 12

Figure 8.3. A. Schematic cross section of Ridge Basin, summarizing essential stratigraphic relationships, from Crowell and Link (1982). B. Sedimentary-tectonic chronology of Ridge Basin and associated finats, from Crowell (1982). 367 at 4.5-5.0 Ma on the basis of its vertebrate fauna, was displaced -220 km from its source area inboard of the San Andreas fault, or close to the full San Andreas offset of 240 km in this area. In summary, the San Gabriel fault is now thought to have started moving between 12 Ma and 8 Ma; slip was transferred inboard to the San Andreas fault at -5 Ma, at which time northwesterly translation of the outboard block accelerated (Figure 8.4). Initiation of the fault system therefore corresponds approximately with the period of mafic igneous activity around the mouth of the Gulf (13-8 Ma, peak -10 Ma) and the first marine incursion in the Gulf of California (-12 Ma); the transfer of slip from San Gabriel to San Andreas and acceleration of the fault system correspond approximately to the beginning of sea-floor spreading in the Gulf mouth (5-4 Ma).

Matti et al. (1985) and Frizzell, Mattinson, and Matti (1986) cast some doubt on this scenario by the discovery of part of a small, distinctive Triassic pluton between the San Andreas and San Gabriel faults which is offset from its matching half inboard of the San

Andreas fault by only 160 km. This well-documented piercing point appears to contradict the 240 km offset interpreted by earlier workers.

Matti et al. (1985) explained this discrepancy by postulating that the missing strike slip was accommodated on the San Francisquito fault

Within the San Gabriel range. Meisling and Weldon (1986) suggested that the offset portion of the pluton was thrust across the San Andreas after it became active, so that the 160 km offset might not record the entire displacement on the San Andreas. 368

350 Pacific/North American plates (Atwater a Molnar, 1973)

Baja California (Moore & Curray,1982) 300 Southern California (Croweft7 1982) n g o Go ,o ve 250 ,a03, o a.

• • 200 2

150

1 100

a 07 0 50

3— late Quaternary slip rates (Weldon, 1984, Weldon & Sieh, 1985, Sieh & Jahns, 1984)

Figure 8.4. Time-displacement diagrams for parts of Gulf of California-San Andreas transform system. (1) Pacific-North American relative finite plate motions, based on Atwater and Molnar (1973). (2) Motion of Baja California relative to North American plate, based on model of Moore and Curray (1982). (3) Motion of San Andreas fault system, southern California, in vicinity of San Gabriel Mountains, after Crowell (1982). (4) San Andreas fault system in central California, after Clarke and Nilsen (1973). (5) Late Quaternary slip rates inferred from geomorphic evidence; note discrepancy with Pacific-North American plate motion. 369 Tighter and less equivocal constraints on the total offset across the San Andreas have been found farther to the north, most notably the Neenach and Pinnacles volcanics (24.5 Ma) which are offset by -315 km (Matthews 1976). Clarke and Nilsen (1973) summarized the

history of the San Andreas fault in central California in a time-

displacement diagram (fig. 8.4). However, because of the complex history of the Transverse Ranges, evidence from north of there may not be applicable farther south (Hornafius 1986), such as in the Salton Trough region.

Geomorphic evidence, primarily offset fluvial terraces with

radiocarbon dates, has been found for late Quaternary slip rates of the

San Andreas fault. In the vicinity of the San Gabriel block, Weldon

(1984) documented 35 + 5 mm/a of right slip on the San Andreas, and

later refined that figure to 37.5 + 5 mm/a. These results are similar

to rates of 34 + 3 mm/a for the last 3.7 ka and 36 + 5 mm/a for the last 13.2 ka, obtained by by Sieh and Jahns (1984) at Wallace Creek, north of the Transverse Ranges.

San Gorgonio Pass region

In San Gorgonio Pass the San Andreas fault splits into three main branches, the San Jacinto, Banning, and Mission Creek faults, and

Possibly other minor strands (Matti et al. 1985). Total displacement is well determined for the San Jacinto fault (Sharp 1967) but not for the other two major faults. Allen (1957) estimated a minimum displacement of 7 km on the Banning fault, based on separation of the late Pliocene or Pleistocene San Timoteo Formation, which contains distinctive 370 piedmontite-bearing clasts, from its probable source area. However, this offset apparently pertains only to part of the complex Banning fault zone, and should be regarded as a minimum between Block 3 and more outboard blocks.

The Mission Creek branch has probably accommodated much larger

displacement, based on clast compositions at the Whitewater Canyon

locality. Peterson (1975) described distinctive clast types in the

Miocene Coachella fanglomerate, including potassium-feldspar porphyr-

itic quartz monzonite, and boulders of magnetite. He identified a

possible source area in the Cargo Muchacho Mountains near Yuma, which

would imply an offset of -215 km. Dillon (1975), in a study of the

southern Chocolate Mountains, discovered a possible alternative source

area, which would require only 180 + 20 km of offset. Matti et al.

(1985) have proposed a scenario for tectonic displacements for the San

Gorgonio Pass region; to obtain an internally consistent history they

used only 160 km total offset on the Mission Creek branch, which is

within the estimated error of Dillon's (1975) interpretation.

San Jacinto fault

Total displacement on the San Jacinto fault system is con-

strained by -24 km total offset of the Santa Rosa cataclasite zone

(SRCZ) and various plutonic contacts (Sharp 1967). Bartholomew (1970)

presented a slightly different geometric analysis and obtained a total

offset of 28 km. Late Quaternary slip rates have been estimated at 8-12

minia (minimum) by Sharp (1981), 6-13 mm/a by Prentice, Weldon, and Sieh

(1986), and 9 + 1 mm/a (minimum) by Rockwell, Merrifield, and Loughman 371 (1986). Savage et al. (1979) reported a four-year average deep slip rate of 17 + 9 mm/a for the San Jacinto fault based on geodetic trilateration.

As the San Jacinto fault approaches the Salton Trough, it splits into two main branches which have been variously named (Dibblee 1954, 1984, Sharp 1967, Bartholomew 1970); they will be referred to

here as the Clark branch (to the north) and the Coyote Creek branch (to

the south). The Clark branch is the more direct extension of the San

Jacinto, but appears to disappear under intensely deformed Neogene

strata with dominantly east-west fold axes (Plate 2; Dibblee 1984). The

Coyote Creek branch is connected to the San Jacinto fault by a series

of small right-offset pull-aparts (Sharp 1975); it is throughgoing and cuts Neogene and Quaternary strata (Plate 2).

Farther to the southeast the Coyote Creek branch appears to

split again into the Superstition Hills and Superstition Mountain

faults, although geometric details are obscured by Quaternary alluvium

(Plate 2). The "San Jacinto fault" identified by Merriam (1965) in

Sonora on the eastern margin of the Colorado delta plain is probably an

offset extension of the San Andreas and Imperial faults rather than a

true continuation of the San Jacinto fault; Merriam's San Jacinto fault

is referred to in this study as the Cerro Prieto fault (Figures 1.1, 1.2, Plates 1, 2).

To the southeast of the area studied by Sharp (1967) and

Bartholomew (1980), in the western Salton Trough, some of the transcurrent motion of the San Jacinto fault may be accommodated by 372 distributed shear strain. Dibblee (1984), Sharp and Clark (1972), Dronyk (1977), and Wagoner (1977) mapped extensive folding indicative of north-south contraction in Neogene strata of the Borrego Badlands, San Felipe Hills, Ocotillo Badlands, and Superstition Hills (Plate

2). The dominant trend of fold axes is deflected near the traces of the

Clark branch (projected) and Coyote Creek branch of the San Jacinto fault system. Bogen and Seeber (1986) reported paleomagnetic evidence for 20-30° of clockwise rotation of the Pleistocene Ocotillo Conglom- erate in a small pull-apart basin along the Coyote Creek branch. In summary, the mode of shear-strain accommodation in the northwest is essentially by strike-slip offset of two rigid blocks, but toward the southeast it changes to distributed shear strain accommodated by multiple strike-slip faults, north-south shortening, and clockwise rotations.

Timing of slip on the San Jacinto fault is not tightly constrained. Sharp (1967) postulated that motion was initiated since the beginning of the Pliocene; Bartholomew (1970) suggested late Pliocene.

As discussed in Chapter 6, other evidence for tectonic activity in this area is from sandstones of the late Pliocene-Pleistocene (?) Borrego and Brawley Formations in the San Felipe Hills, which contain Colorado

River-derived sandstone evidently reworked from the west (Wagoner

1977), thus implying uplift and erosion of Neogene basin-fill on the western magin of the trough (Chapter 6). 373 Elsinore fault

Total displacement on the Elsinore is not as well established as for the San Jacinto fault due to a paucity of well-defined piercing points. Lamar and Rockwell (1986) recently reviewed the extensive literature on this fault; a more cursory treatment will be presented here. The largest postulated displacement is -40 km, based on offset of an inferred Paleocene shoreline of the Los Angeles Basin (Sage 1973, Crowell and Ramirez 1979, Link, Squires, and Colburn 1984). However, in

other areas, displacements of only -10 km or less can be demonstrated (Lamar and Rockwell 1986). Crowell and Sylvester (1979a) estimated -6

km based on the size of the Lake Elsinore pull-apart basin. Kennedy

(1977) reported a minimum of 5 km based on offset of a Pleistocene

(-0.7 Ma) facies boundary in the vicinity of Chaney Hill. Lowman (1980)

has argued for essentially no strike-slip offset on a large part of the

Elsinore fault system, based in part on continuity of metamorphic

foliation across a mapped portion of the fault at Campbell Grade. How-

ever, Pinault (1984) inferred that the observed foliation occurs in a

large, coherent landslide mass that crosses the fault. Estimated slip rates on the Elsinore Fault include a minimum of 7 mm/a (Kennedy 1977),

3-5 mm/a (Pinault 1984; Pinault and Rockwell 1984), 5-6 mm/a (Millman and Rockwell 1986), and 3-'-6 mm/a (Vaughan and Rockwell 1986); Savage et al. (1979) reported a four-year average deep slip rate of 7 + 9 mm/a for the Elsinore fault, based on geodetic trilateration.

Within the Elsinore fault system, particularly on the western flank of the Salton Trough, right slip may be regionally distributed, 374 as with the San Jacinto fault. The Elsinore may splay out into the Agua Caliente, Earthquake Valley, and other small faults for which strike- slip displacement has not been demonstrated. Johnson et al. (1983) have estimated -35° of post-depositional (i. e. post-0.9 Ma) rotation of the FCV section based on discordance of paleomagnetic declinations in the

Plio-Pleistocene section (Figure 8.5). Johnson et al. (1983) estimated that this rotation would require -7 km of distributed dextral shear, for a rate of -8 mm/a; in Chapter 6, it was estimated that dextral shear could be as much as 12 km. Mace (1981) observed a wide dispersion of paleomagnetic declinations in basalts of the Miocene Alverson Forma- tion in the eastern Coyote Mountains and the Carrizo Impact area. These areas are complexly faulted, particularly the eastern Coyote Mountains, suggesting that the inferred rotations could have been accommodated within small structural domains as discussed by Wells and Coe (1985).

Mapping in the Sweeney Pass area in the present study suggested 3 + 1 km of offset of the Diablo-011a-Hueso-Canebrake facies tract across the

Elsinore fault (Plate 4). By adding this figure to the 7 km associatied with rotation of the FCV block, at least 10 km of post-mid-Pliocene right slip can be accounted for. Pinault and Rockwell (1984) reported a slip rate of 4 + 1 mm/a based on geomorphic evidence along the southern flank of the Coyote Mountains; combined with 8 mm/a associated with roation of the FCV section, a total rate of -12 mm/a is possible, greater than rates considered in recent tectonic models (Weldon and

Humphreys 1986, Bird and Rosenstock 1984). 375

alpha - 95 NORMAL SITES alpha-95 REVERSED SITES

Figure 8.5. Stereonet projection of virtual geomagnetic poles for 0 PCV section, from Johnson et al. (1983). Approximately 35 of post— depositional clockwise rotation is implied. 376 Alternatively, rotation of the FCV block may offset slip from the Coyote Creek fault to the southern Elsinore-Laguna Salade fault, in which case the slip rate on the northern Elsinore fault may be less.

The distribution of epicenters of magnitude >6 earthquakes suggests that the Laguna Salada fault is more active seismically than the northern Elsinore fault (Collette and Ortlieb 1984), which tends to support this interpretation.

As with the San Jacinto fault, timing is not well constrained.

Most of the evidence discussed so far is for post-mid-Pliocene or Quaternary displacement (Chapter 6). No evidence of syn-tectonic

conglomerates older than Pleistocene has been reported from along the

Elsinore fault, even in the Salton Trough area.

Algodones fault

The Algodones fault represents a major down-to-northeast offset

on basement of the Yuma desert based on well data and gravity and

magnetic surveys; this fault separates the San Luis and Fortuna Basins

(Olmsted et al. 1973). This fault is approximately aligned with the San

Andreas fault farther to the northwest, and has been suggested as a

Possible former extension of the San Andreas (Crowell and Ramirez

1979). However, no strike-slip displacement has been conclusively

demonstrated. Short, north-south, en echelon surface faults in northern

East Mesa, on strike between the San Andreas and Algodones faults, may

indicate recent right slip on the Algodones fault system (Heath 1980). 377

Salton Trough region

As the San Andreas fault is traced toward the southeast, evidence for active slip dies out past the Saltoil Sea. Geodetic trilateration surveys of the Salton Trough demonstrated that right slip along the San Andreas is stepped right to the Imperial fault. The zone of right step is characterized by microseismicity in roughly rhomb- shaped areas (Johnson and Hill 1982, Fuis and Kohler 1984). Savage et al. (1979) reported four-year average deep slip rates of 25 + 8 mm/a

for the San Andreas fault adjacent to the Salton Sea, and 47 + 21 mm/a for the Imperial fault, both based on geodetic trilateration. While the

error bars on these data overlap, the higher estimated rate for the

Imperial fault suggests that it may accommodate some of the slip from

the San Jacinto fault system as well. Studies of faults in and around

the Cerro Prieto Field indicate that this geothermal field occupies a

pull-apart basin which offsets slip from the Imperial fault to the

Cerro Prieto Field (Vonder Haar and Howard 1981).

Ague Blanca fault

This fault in northern Baja California was described by Allen

et al. (1960), who demonstrated a minimum of -10 km of right slip and

possibly as much as 20 km, based on offset slivers of distinctive

basement lithologies in the fault zone. Hatch and Rockwell (1986) found a rate of geomorphic evidence for late Quaternary slip, and estimated faults have 2-5 mm/a. South of the Agua Blanca fault, no throughgoing Quaternary been found in Baja California with evidence for Neogene or

displacement. 378 Displacements vs. rates If the reference model were correct, Quaternary slip rates would be proportional to total displacement for individual faults and for the entire transform system. On a plot of displacements vs. rates, all components of the system would plot on the 5.5 Ma isochron (Figure 8.6). This prediction can be tested with-the data just summarized. In

Figure 8.6, data for the San Andreas fault in the Salton Trough region

plot above the 5.5 Ma isochron, suggesting that this component of the

system has been moving disproportionately slowly in the late Quater-

nary. Data for the San Jacinto fault plot below the same isochron,

suggesting that late Quaternary movement has been disproportionately

slow. Data for the Elsinore and Agua Blanca faults are more ambiguous

due to the large proportionate errors, but plot more below the isochron than above.

These results suggest that the San Andreas fault south of the

San Gorgonio Pass has decelerated relative to the entire transform

system as that system evolved, while the outboard faults (San Jacinto,

Elsinore, Agua Blanca) accelerated. This conclusion is reasonable in

view of the geometry of the San Gorgonio constraining bend (Plate 1);

outboard transfer of slip south of the constraining bend would tend to

straighten the transform boundary and thereby reduce its internal

resistance. It also indicates that the reference model is overly simplistic and that possible deviations from the reference model should be considered. 379

RATES: SPREADING RATES AT MOUTH OF GULF — (Curray et al., 1982; E 300 Ness et al., 1981) a PACIFIC / NORTH AMERICAN PLATE MOTION (Minster & Jordan, 1978)

Ui LATE QUATERNARY 2 GEOLOGIC SLIP RATES LU (Weldon & Sieh, 1985; o Sharp, 1981; < 200 Johnson et al., 1983; Pinault & Rockwell, 1984; Q. Hatch & Rockwell, 1986) o GEODETIC STRAIN RATES (Savage et al., 1979) < DISPLACEMENTS: 0 100 Allen et al., 1960 Matti et al., 1985 Bartholomew, 1970 Moore & Curray, 1982 Crowell & Peterson, 1975 Sylvester, 1979 Rockwell & Dillon, 1975 Sylvester, 1979 Johnson et al., 1983 Sage, 1973 Matthews, 1976 Sharp, 1967

Figure 8.6. Total displacements vs. rates for components of the San Andreas fault system in the Salton Trough. Diagonal lines are isochrons representing the time required to attain a given displacement at a given rate. Components plotting below the reference model isochron are thought to be accelerating (or have accelerated) relative to the entire System, above the line to be decelerating (or have decelerated). 380 Analysis of displacement histories Because of the unavoidable errors in tectonic reconstructions this sort, an of attempt was made to define envelopes of uncertainty to encompass the range of geologically reasonable displacement histories for the major tectonic blocks in the study area (Figures 8.7-8.11). In all plots, displacement is relative to the NAP (Block 1) and relative to the present position in a southeasterly direction. For Block 2 (San

Luis Basin) it was not possible to construct such a plot due to lack of direct evidence for strike-slip offset on the Algodones fault. The

Salton Trough region is treated as seven rigid blocks, as in the reference model, although rotation of the FCV block as a possible contributor to total slip on the Elsinore fault is taken into account.

Only right-lateral displacements are considered, so that no individual displacement rates (relative to the NAP) can be greater than that of the Pacific Plate (Figure 8.4). Erratic motions involving long-term stick-slip behavior or very unsteady slip rates are not considered.

Block 7 (Figure 8.7)

The displacement envelope for Baja California was constructed to honor two main constraints: magnetic anomalies in the mouth of the

Gulf of California for the last -4 Ma, and a minimum total displacement of -300 km, based on the combined displacement on the San Andreas and

San Gabriel faults. The lower limit on the displacement envelope was drawn to accommodate the inference that the San Gabriel fault began to move as early as 12 Ma (Figure 8.4). The total San Andreas displacement 381

380 \ I \ \\\\\\\\\\\ \ additional 80 km possible Block 7 \ \ on Agua Blanca + Elsinore faults Il litlICHEI111111 3110111110111111:11 I 300,- 3.15±5 km, Neenach & es . Pinnacles volcanics -- (Matthews, 1976) C o &a-€-1-- 260 movement on San Gabriel fault o n. as early as 12 Ma t (Crowd, 1982) a a o 200 it beginning of magnetic anomalies, 2 mouth of Gulf m (Currey et 81.1982) 2... 160 co I rc 58±3 mm/a 1 100 (Minster & Jordan, 1978) 0 a

"Eta reference model 5 (after Moore & Curray,1982)

Displacement history with uncertainty envelope Figure 8.7. shown (diagonal for tectonic block 7. Locations of blocks are ruling) at 12, 7, in Figures 1.1 and 1.2. Solid circles and vertical error bars 5.5, 4, and 3 Ma represent reference-model displacements and uncertainties plotted on palinspastic base maps (Figures 8.12-8.15) 382

upper bound of Block 7 envelope, less lower bound of Agua Blanca envelope Block

300 o 11 250 o Q. lower bound of Block 7 envelope, less upper bound of Agua Blanca envelope •ID 200 2 iso 2to cc reference model i 100 ▪ geodetic slip rate (Savage et 81,1979) Tt

Ague Blanca fault:10 - 20 km (Allen et al, 1960)

Figure 8.8. Displacement histories with uncertainty envelopes for Agua Blanca fault and tectonic block 6. 383

350 ... Block 6 316t5 km, Neenach - Pinnacles voicanics (Matthews, 1976) 171n11s111111n1101M1 EMIL 300 upper bound for Block 6, less o lower bound for 250 o Elsinore fault o. IS

• 200 lower bound for Block 6, less upper bound for Elsinore fault

rti> 150

reference model a) E 100 a -Noll geodetic slip rate a (Savage et a1,1979) a o 50

Figure 8.9. Displacement histories with uncertainty envelopes for Elsinore fault and tectonic block 5. 384

350 _ Block 4

upper bound Block 5, less lower bound for San Jacinto fault 300 o

o 200 1200 2 lower bound Block 8, 150 less upper bound for San Jacinto fault

reference model

geodetic slip rates (Savage et 81,1979) a 0 50 San Jacinto fault - 26 km (Sharp. 1987) to 30 km (Bartholomew, 1970)

14 18 18

Figure 8.10. Displacement histories with uncertainty envelopes for tectonic block 4. 385

Block 3

215±20 km (Peterson, 1975)

180120 km (DMon,1975)

reference model

-41111 geodetic slip rate (Savage et a1,1979)

Figure 8.11. Displacement histories with uncertainty envelopes for tectonic block 3. 386 could be closer to the 315 km offset on the Pinnacles-Neenach volcan- this figure ics, but is less directly applicable to the Salton Trough because of the intervening Transverse Ranges. Slip accommodated by the Elsinore (40 km) and Agua Blanca (20 km) faults could be accommodated farther to the north outboard of the San Andreas and San Gabriel faults

(this is not clear on the basis of geometry) so these offsets should be added in to get a reasonable upper limit (-380 km). Whether the Gulf of California could accommodate this much displacement has not been estallished because of insufficient data on crustal extension of the Gulf margins.

Block 6 (Figure 8.8)

The displacement envelope for this block was obtained by estimating a similar envelope for the Agua Blanca fault and subtracting it from the envelope for Block 7. Upper and lower limits of 20 km and

10 km (Allen et al. 1960) are used for the total displacement. Since timing is not known, late Quaternary rates (Hatch and Rockwell 1986) were extrapolated back in time to obtain the rest of the envelope. A slower rate or earlier initiation would have a minor effect on the displacement envelope for Block 6.

Block 5 (Figure 8.9)

A displacment envelope similar to that for the Agua Blanca fault was constructed for the Elsinore fault, using maxima and minima of 40 km and 10 km, as discussed earlier, and extrapolating late Qua- ternary slip rates (7-12 mm/a) back in time. The possibility of two 387 phases of movement, one Plio-Pleistocene and the other associated with Miocene rotations in the Transverse Ranges, can be accommodated within the uncertainty envelope. For Block 5, uncertainty on total displacement is essentially the same as for Block 7; uncertainty for the last 4 Ma is somewhat greater than for Block 7.

Block 4 (Figure 8.10)

A time-displacement envelope was constructed for the San

Jacinto fault similar to those for the Elsinore and Agua Blanca faults,

and subtracted from the envelope for Block 5. If some of the total

offset on the San Jacinto fault took place during Miocene rotation of

the Transverse Ranges, the lower boundary of the envelope would have to

be placed lower. Because of distributed shear in the western Salton

Trough region, various individual localities may have had slightly

different displacement histories. However, the possible error resulting

from distributed shear is less than the error from other sources.

Block 3 (Figure 8.11)

Displacement of this block is constrained more closely relative

to the NAP (Block 1) than to Baja California (Block 7). A total dis-

placement of 180 + 20 km (Dillon 1975) to 215 [+ 20?] km (Peterson,

1975) was used. The timing of this displacement is known only to be

post-10 Ma, based on a dated volcanic flow near the base of the the

Coachella Fanglomerate (Peterson 1975). The uncertainty envelope was constructed on the basis of two end-member possibilities: (1) that all

the early motion between Blocks 1 and 7 was accommodated by the Mission 388 Creek fault, which gives the lower boundary on the envelope, and (2) that the Mission Creek fault became recently active and has since accommodated essentially all of the motion between Blocks 1 and 7. This second possibility is quite unlikely in view of the high late Quater- nary slip rates on the San Jacinto fault, but there are no other obvious constraints on an upper boundary for the envelope.

Effects of additional distributed shear

Additional unrecognized distributed shear would affect the

validity of these displacement-history envelopes. Additional shear strain between the mouth of the Gulf and the Salton Sea would have the effect of "lowering" the displacement curves for localities in Blocks 7 through 4. Shear strain accommodated inboard of the San Andreas fault would "raise" the displacement curve for Block 3. More importantly, it would mean that the Chocolate, Laguna, Cargo Muchacho, and Gila Moun- tains could not be regarded as rigidly attached to the North American

Plate. This would in turn have implications for the position of the

Colorado River entry point relative to the western margin of the Salton

Trough (Winker and Kidwell 1986).

This possibility must be taken seriously because of the problem of late Quaternary slip-rate budgets. Triangulation across the Salton

Trough shows that historic shear strain can only account for -40 mm/a of motion between the Pacific and North American Plates (Savage 1983).

Even if an additional 2-5 mm/a on the Agua Blanca fault (Hatch and

Rockwell 1986) is taken into account, a shortfall remains between total measured strain rate and the -52-56 mm/a for the mouth of the Gulf of 389 California (Figure 8.2) or 56 + 3 mm/a between the Pacific and North

American Plates (Minster and Jordan 1978). Baja California appears to have behaved as a rigid block during the late Cenozoic, based on the lack of throughgoing faults south of the Agua Blanca (Plate 1). There- fore, any shear-strain deficit is apparently taken up inboard of the

San Andreas fault. Sauber and Thacher (1984) estimated -5 mm/a of right slip in the Mojave Desert between the Helendale and Camp Rock faults based on triangulation and trilateration studies between 1934 and 1982

(see King 1985 for a more conservative interpretation). Thompson and

Burke (1974) estimated an overall extension rate of 7 mm/a in an east- west to southeast-northwest direction across the entire Basin-and-Range province. These rates may be nearly sufficient to account for the slip-rate deficit.

A remaining question is whether this slip-rate deficit applies to the entire late Cenozoic history. While the present San Andreas fault may accommodate only about two-thirds of the total interplate slip, the San Andreas (including the San Gabriel) north of San Gorgonio

Pass has accommodated -300 km of right slip, which is almost certainly more than two-thirds of the total opening of the Gulf of California.

This suggests that transfer of slip from the San Andreas to other outboard and inboard faults and distributed shear may have developed

Progressively as the system evolved. 390 Trajectories

In the reference model, azimuths of 315° were assumed for all trajectories, based on the dominant trend of strike-slip faults in the

Salton Trough area. However, according to Blake et al. (1978), the

finite-motion azimuth of Pacific Plate motion relative to North America

at the latitute of the Salton Trough was 319° for the last 4.5 Ma. The

discrepancy between plate motion and fault orientations may be

explained by tectonic models proposed for southern California by

Garfunkel (1974) and Weldon and Humphreys (1986). Garfunkel (1974)

suggested the westward translation of the Sierra Nevada region relative

to rigid North America, a result of Basin-Range extension, could

account for both the left-lateral displacement on the Garlock fault and

the "big bend" of the San Andreas fault (Plate 1). In his model, this

portion of the San Andreas fault has undergone counterclockwise rotation.

Similarly, Weldon and Humphreys (1984) suggested that southern

California outboard of the San Andreas fault may be moving obliquely

toward the Pacific Plate and undergoing slow counterclockwise rotation

(Plate 1); otherwise a severe room problem would result at the Trans- verse Ranges. This model would imply that faults in outboard southern

California are also being rotated counterclockwise, so that their present orientation may not be representative of long-term azimuths of tectonic displacements. For this reason, an azimuth of 320°, closer to the relative Pacific-North American plate motion, is regarded as plaus- ible for tectonic trajectories in the Salton Trough. 391

Paleogeographic maps

Palinspastic base

Using the preceding analysis, error bars were added to restored

positions of individual localities based on the reference model

(Figures 8.12, 8.13). Uncertainties in the northwest-southeast direc-

tion are based on envelopes of uncertainty constructed in time-

displacement diagrams (Figures 8.7-8.11). In the southwest-northeast

direction (azimuthal errors), bold error bars indicate the result of

320° azimuths rather than 315°; light error bars show the effect of an

additional 50 (an arbitrary figure) of azimuthal error.

The largest errors in the northwest-southeast direction occur

in the late Miocene, due to uncertainties in total displacement and the

time of initiation for the San Andreas-Gulf of California transform

system. Uncertainties associated with individual localities are not mutually independent. For example, the positions of the FCV and San

Felipe Hills localities relative to each other are much more tightly

constrained than the lengths of their error bars. Therefore, more local

palinspastic reconstructions, such as the one developed in Chapter 6

(Figure 6.16), may have smaller uncertainties than the error bars would indicate.

Faults and spreading centers

To a first approximation, positions of spreading centers can be reconstructed on the assumption that they behaved like oceanic spreading centers, in that the centers migrate at the half rate 392

Ce) 9-4 o CO Co 0 • o. o o o 44k •ri (1) o Co

• 1-1 a) cj

• Ct C.) r—I • a4 4-) X a) .4 0) 0

(1) 0 .0 0 ai

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0 CO tH 1-n CI, 0)

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2 cv o + •

393

Ce; >.• 0 El tO 0 • U

C • in -ci 1.4 0 0 • 11 • cn o o b0 g •ri ostv

04 02 • tr: 0 a I -0 • ,rz 02 0 0 4-) 0 XI 04

4) •4 1.4 O g 0

c• cri 0 -H 0 o - O 1:14 • 04 ci) r•-.1 • ri 0 04 -H g 0)0 O• 0 4:2 c4.4 C •ri (.1 a,0) 4-1 .1-1 O 'H g O g

g• 0 C12 • -P O •4) 4) 4.) 0 rn '00 O 0 0 40 10 0.) • 0 4-) •,•-i 0 •-t O U A, 0 • • - 0 "0 • C.) 0 CO ' H 4)

0 g 0, 5.1 • o. r.., a, 43) • 4-) 0:1 Lo 394 relative to either plate. However, if this procedure is carried back to the beginning of transform motion, serious room problems are encountered. This suggests that the spreading-center geometry has changed with time. This possibility has previously been suggested, based on (1) heat-flow modeling of the Salton Trough by Lachenbuch, Sass, and

Galanis (1985), who suggested that the spreading center configuration is short-lived relative to the age of the trough; (2) high heat-flow mummdies in the Salton Trough not associated with historic microseismicity, possibly representing extinct spreading centers; and (3) bathymetry around Isla Angel de la Guarda (Bischoff and Henyey 1974) which suggests the existence of pull-apart basins on both the northwest and southeast sides of the island. Sharman (1976) has suggested that the Carmen Basin in the Gulf of California (Plate 1) evolved by a succession of spreading-center jumps. Therefore, the reconstructed tectonic maps (Figures 8.12, 8.13) include speculative spreading centers (stippled).

The Elsinore, San Jacinto, and Agua Blanca faults are shown

only for the 3 Ma reconstruction and present base map because of

evidence that they became active in the late Pliocene or Pleistocene.

Marine transgression and regression

The distribution of stratigraphic data is inadequate to

ascertain the precise location of shorelines of the Neogene Gulf of

California, particularly for the eastern margin of the trough.

Nevertheless, paleogeographic maps are a useful way to illustrate the

transgressive history of the nascent Gulf of California and the 395 initiation of the Colorado delta. These maps were drawn on the palinspastic base maps derived from the reference model; possible deviations from the model are indicated by the error bars (Figures 8.14, 8.15).

Timing of marine transgression and regression are constrained by geochronologic data summarized in Figures 7.1 and 7.9. Marine water reached Isla Tibur n by 12 Ma (Smith et al. 1985), San Felipe by 6 Ma

(Boehm 1982, 1984), Whitewater Canyon before 6 Ma (Matti et al. 1985), and the Bouse Embayment by 5.5 Ma (Shafiqullah et al. 1980). Late

Miocene and early Pliocene shorelines in the San Felipe area are based on maps in Boehm (1982). The Bouse shoreline is based on outcrop and subcrop distribution and present distribution of topographic lows.

Regression from the Bouse Embayment prior to 3.8 Ma

(Shafiqullah et al. 1980) may have been caused in part by deltaic progradation, although the author did not find a progradational deltaic sequence exposed in the Bouse Embayment. Epeirogenic uplift of the

Bouse Embayment (Lucchitta 1972, 1979) was probably the more funda- mental cause of the Bouse regression. Marine regression from the

Whitewater Canyon area prior to 6 Ma (Matti et al. 1985) was caused by progradation of locally-derived clastics. Regression from the San

Felipe and Isla Tibur n localities probably resulted from tectonic uplift. Regression from the "Hermosillo Basin" may have been caused by progradation of a local delta. Regression from most of the Salton

Trough resulted from progradation of the ancestral Colorado delta;

Colorado River-derived sediment first reached the Trough by 4.3 396

o o o eg o

o o _ o

o o

+ \\\\'‘ \ \\\\\‘``') \\ 2 +- \+ E e0 w

C\ +

397

o 4.. CO (4-1

/ l'2 0) C .0 0 a -1. 03 co03 ^10 03 o + g A co. 1 zo 4 . • I 0 0E. -1 12 0:, .65 • a D2' 5 1 - ..... • 2 T. 4. ›..) 0 CL4 74 LL, -=> g.. , oi-t 4 • . 3 8 g 6+ • 0 0 10 04 0 b0 g H re

O 4.4 c..1 a) O a) O -0 •n••( - n—n • 0 a. (1) O 0 0 44-4 cp 0 0) O .0 Ow 0

• g u • G) -V 04 0 E .00 C4-4 ai O 0 Le (3.) be -P 0 Q) -P al X $.4 fr.a 0 • • 0 4r)

03 0 a) 0) 'DO 1--4 z C) .00 b0 • n••1 $.4 0 (1) ^V 0 O - E a4 398 Ma, and the delta plain spanned the Trough and had prograded as far south as the FCV section by 4.0 Ma.

Tectonic-paleogeographic history Plate-tectonic context

Many aspects of the late Cenozoic history of southwestern North

America have been successfully explained in terms of interaction of the

North American, Pacific, and Farallon Plates, as first proposed by

Atwater (1970). For example, Snyder, Dickinson, and Silberman (1976)

demonstrated that the shutoff of arc-related volcanism in California

and Nevada from 20 Ma to the present tracks the projected trend of the

restored Mendocino Fracture Zone (Figure 8.16). More recently, Glazner

and Supplee (1982) demonstrated that a locus of intense volcanism

migrated northward along with the projected trend of the Mendocino

Fracture Zone. Dickinson and Snyder (1979a) developed the "slab-window"

model based on interaction of three original plates to explain late

Cenozoic changes in volcanic geochemistry in southwestern North America.

Ultimately, then, it would be desirable to relate the history

of the Salton Trough to the same model of plate interactions. As discussed in Chapter 7, the onset of circum-Gulf rifting can be dated approximately by changes in volcanic geochemistry (Figure 7.9). How- ever, the initiation of rift-related volcanism (Figure 8.16) correlates

Poorly with the predicted migration of the Rivera Triple Junction

(Atwater 1970) or the slab-window model' of Dickinson and Snyder (1979a). • • •

399

tr o

;

C,) vavr N313 n11:10 -vivavno — • VII1100 831N384800 08'00 :in;1$;i111q1,

05 1-4 4-) • e-.4 0) .+J 0 0 O 0 -H S-1 O -M..) 0 *S.Lw X II 3.1.v1030N0 • t1-( X CU 0 U L . -1•J 0018 3Avr0vi • 04 0 0 o a., a., 4, a) N (1J O 0) a) -E 0.) -P 4.) a) .0 C.) U) Q4-4-> OOSIONVIdd 0) 0) 4)4-a 0 >AIS 4) •H la. -P O 3 .4.) F-1 • •H •ri g O 0 g 0 0 0 ONIOOCINSIN 0 "CI "ci •,•-n o ▪O c.) a) D a) cz 4.> 04) 5. 0 -P 0 12.) Cl) 31.U.V3S- CO M• •n D G 0+ 1-1 V) a) u. 2 To O C1..-0 ONV1SI (Atwater, 1970) ."- O Cl) 4-0 n =No de 0 =11; 0 83A1103NVA .00,00,„„...... C‘.11.1“. PC,11•••••• bo 4.) X(.) O 14 • 0r0,..< M.:1 0 0 .0 LU - E (3) c4-1 4-) a. •r4 U) •,-1 La-0 Pdf; Et 0 1 3 FE 0 I- Cr) ..... • Q) 1.N3dON CO O 30NIbld- . : •,-; ci-, • - o to In Cf CO 0 0 -0 y- (1) 0. a) 0 • .r.1 0.) • 1.. C./ a) CD .0 4 -P •ml L 0 Cul 0 0 0 • LO U) -p b0 • 0 ^c) O d) g .0 Z -P 400 An explanation for this serious discrepancy may be found in the

complex history of microplate motions off Baja California. Chase et

al. (1970) recognized that a microplate separated from the Cocos Plate

and began to spread normal to the Baja California trench, rather than

oblique to the trench. More recently, Ness et al. (1981), Klitgord and

Mammerickx (1982) and Mammerickx and Klitgord (1982) identified pos-

sible extinct spreading centers off Baja California. Ness et al. (1981)

placed the termination of spreading at 13.7 Ma, whereas Klitgord and

Mammerickx placed it at 12.5 Ma. In either case, this plate reorganiza-

tion would have initiated a transform boundary between the Pacific and

North American plates outboard of Baja California, presumably repre-

sented by the Tosco-Abreojos fault zone (Spencer and Normark 1979).

This plate reorganization can also explain the near-synchronous change

from arc-related to rift-related volcanism in Baja California Sur

(Sawlan and Smith 1984, Rausback 1984). The timing of plate reorganiza-

tion also agrees well with the earliest marine inundation of the Gulf

(Smith et al. 1985).

In contrast to the north-migrating Mendocino Triple Junction,

the corresponding south-migrating intersection of the East Pacific Rise with the western margin of North America probably cannot be modeled as

a simple three-plate geometry with a discrete triple junction. Instead, microplates evidently become sequentially detached from the Cocos

Plate and subsequently transferred to either the Pacific Plate or North

American Plate. This has happened offshore from Baja California (Chase et al. 1970, Ness et al. 1981, Mammerickx and Klitgord 1982), within 401 the Gulf mouth (Rivera Plate), and may be now in progress south of Cabo Corrientes (Luhr et al. 1985). As Dickinson and Snyder (1979b) pointed out, a ridge-trench-transform triple junction of the Rivera type

involves subduction of very young and hot oceanic crust, which may

resist subduction and therefore may not behave in a geometrically

simple fashion. Menard (1978) explained fragmentation of the Farallon

Plate as a response to the strong trench-wise gradient in the age of the subducting crust and resulting gradient in the magnitude of the "slab-pull" force.

Proto-Gulf rifting

From 14 to 12 Ma, much of the circum-Gulf region underwent

major changes, including widespread termination of subduction-related

volcanism, initiation of rifting, and marine inundation as far north as

Isla Tibur n (Figure 8.16). These events appear to be direct or

indirect consequences of the extinction of sea-floor spreading and

initiation of a transform plate boundary offshore from Baja California.

In the Salton Trough, however, there was evidently a consider-

able lag between the onset of rifting and marine transgression (Figure

8.16). Rift-related volcanism began as early as 22 Ma, probably a

response to the initial encounter of the Pacific Plate with the North

American plate at -30 Ma (Atwater 1970), but delayed due to the slab- window effect (Dickinson and Snyder 1979a). However, no marine connection developed to the north, and circum-Gulf rifting had not yet begun farther to the south. By 12 Ma, the rift extended at least to the San

Jos del Cabo Trough (the probable proto-Gulf mouth) and permitted the 402 first marine incursion. Marine waters probably reached the Salton

Trough region somewhat later, but the timing is poorly constrained.

Transform plate boundary

During the proto-Gulf extensional phase, relative motion

between the Pacific and North American plates was presumably accom-

modated outboard of Baja California (Yeats and Haq 1981) on offshore

fault zones such as the Tosco-Abreojo (Spencer and Normark 1979). A

transtensional regime was initiated in the Gulf of California and

Salton Trough by inboard transfer of slip, probably utilizing the

active continental rift as a pre-existing zone of lithospheric weakness

(Karig and Jensky 1972). The change from proto-Gulf rifting to the

modern transtensional regime may have involved a transitional phase

during which Baja California was only partly coupled to the Pacific

Plate. Evidence for a partially coupled transtensional phase comes from

the displacement history of the San Gabriel fault, which accommodated

displacement between the Baja California block and North America at a

considerably slower rate than the Pacific-North American slip rate (Figure 8.4).

The present transtensional regime, with Baja California essentially coupled to the Pacific Plate, began -5-4 Ma (Figure 8.16), based on initiation of the San Andreas fault adjacent to the San Gabriel

Mountains (Figure 8.3); sea-floor spreading at the same rate as the

Pacific-North American slip rate began by 3.9 Ma (Curray et al. 1982).

Establishment of this fully coupled, transtensional regime is probably what is meant by "opening of the Gulf", a commonly used phrase in the 403 Gulf of California literature. However, marine inundation of the Gulf and Salton Trough preceded this "opening" by a considerable period of time, as much as 7 or 8 million years. This points out the importance

of distinguishing between tectonic history and paleogeographic history when discussing "opening" of the Gulf.

Western Salton Trough

The western side of the Salton Trough experienced a phase of

relative tectonic quiescence between early to middle Miocene rifting

and late Pliocene to Pleistocene transpressive deformation. During this

quiescent period the western Salton Trough behaved rather like a pas-

sive continental margin, similar to margins of the modern Gulf of

California. The subsidence history for the FCV section exhibits expon-

ential decay (Figure 6.4), as expected for subsidence by post-rifting

thermal relaxation, but the early rates of tectonic subsidence were more rapid than would be expected from the Parsons and Sclater (1977)

subsidence curve. The FCV, CIA, and Coyote Mountains sections lack a well-defined "breakup unconformity" (Falvey 1974) such as commonly

occur on trailing continental margins and signal the change from active rifting to passive subsidence (Bally 1981, Scrutton, 1982). Instead, the syn-rift sequences in the present study are characterized by local angular discordances grading laterally into conformable contacts, which may be more characteristic of syntectonic unconformities in areas of rapid net sedimentation (Riba 1976).

Renewed tectonism in the western Salton Trough during the late

Pliocene or Pleistocene is evidently related to outboard transfer of 404 slip from the San Andreas fault to the San Jacinto and Elsinore fault systems. To the northwest of the Trough, this slip was accommodated essentially by two discrete faults, the San Jacinto and Elsinore sensu strictu. Within the western Trough, however, slip was regionally distributed, as indicated by numerous synthetic and antithetic strike-slip faults, block rotation on different scales, and transpressive folding with predominantly north-south contraction (Plate 2). These tectonic distortions complicate the problem of detailed palinspastic restoration

for the western Salton Trough (e. g. Figure 6.16).

Pre-deltaic northern Gulf

Bathymetry of the .early (pre-deltaic) Salton Trough and topo-

graphy of flanking ranges, as inferred from the sedimentary record,

were similar to those of the modern Gulf of California, except that the

early Salton Trough was narrower and probably shallower on average than

the modern Gulf. Alluvial fans developed on the high-relief flanks of

the basin, though not necessarily along fault scarps. Fan deltas with

shallow-water, low-energy fringes, or with local, steep, submarine

aprons deposited by high-density, coarse-grained sediment gravity flows

were the dominant marine depositional systems. Possible tidal influen-

ces on the pre-deltaic basin are difficult to assess, because of the

paucity of physical sedimentary structures due to pervasive biotur-

bation in shallow-marine fan-delta deposits such as the Latrania

member.

Small, closed, bathymetric sub-basins were probably common, and

the Fish Creek Gypsum was probably deposited in such a depression 405 immediately prior to marine inundation. The postulated bathymetric configuration of the pre-deltaic northern Gulf raises the possibility

of local sediment-starved sub-basins in which late Miocene diatomites similar to those at San Felipe (Boehm 1982, 1984), or other organ-

ic-rich sediments, could have accumulated. However, no direct evidence

for such a marine facies in the Salton Trough has been reported.

Colorado River and delta

Uplift of the Bouse Embayment (Lucchitta 1972, 1979,

Shafiqullah et al. 1980) and consequent initiation of the lower Color-

ado River and delta at -5 Ma are closely associated temporally with

development of the fully coupled, transtensional tectonic regime in the

Gulf (Figure 8.16). Whether there is also a causal relationship is not

clear, nor is it apparent why inboard transfer of interplate slip

should be related to a rapid pulse of epeirogenic uplift. In any case,

the large influx of sediment from the Colorado River was responsible

for creating the Salton Trough as a tectonic province distinct from the

Gulf of California.

Observations from the Salton Trough in general and the FCV

section in particular have important implications for the history of

the Colorado River. First, the oldest Colorado River sediment in the

Salton Trough is no older than early Pliocene. This result is con-

sistent with limited subsurface data from the Bouse Embayment which

indicate that no Colorado River sediment underlies the marine Bouse

Formation (Metzger et al. 1973, Metzger and Loeltz 1973), implying that

the lower Colorado River did not exist in pre-Bouse time. Furthermore, 406

if an ancestral Colorado River flowed southward from the Colorado Plateau during the Miocene, as postulated by Hunt (1969), it could not have debouched into the northern Gulf of California, and another sink for Colorado River sediment must be found.

Second, the Colorado River that discharged into the Salton Trough during the early Pliocene was not demonstrably different from

that of the late Quaternary. Petrographic data from Potter (1978) for

late Quaternary Colorado River sand fall within the range of "C"-suite

compositions for Pliocene Colorado River sand (Chapter 2). From a

sedimentological standpoint, paleochannel sandstone sequences in the

Pliocene Diablo member are very similar to paleochannel sands of the

late Quaternary Colorado delta plain as described by Van de Kamp (1973)

(Chapter 6). Although rates of sediment supply to the Salton Trough are

difficult to estimate, the Pliocene Colorado River carried a very large

sediment load, as evidenced by rapid growth of the delta (this chapter)

and high sedimentation rates at the FCV section (Chapter 6). By 4.0 Ma

the delta plain was nearly as extensive as the modern delta plain

(Figure 8.15); subsequently, the delta has been maintained in a predom-

inantly aggradational mode while the Salton Trough widened and deepened beneath it. These results suggest that the modern configuration of the

Colorado River system was created quite rapidly in the early Pliocene, and has not undergone substantial reorganization since then.

Temporary departures from this steady state may have occurred during the late Pliocene and/or Pleistocene. Ingle (1982) found evidence of Pleistocene marine incursions over the delta plain (Chapter 407 7). These could easily result from rapid glacio-eustatic sea-level fluctuations on the order of 100 m, as the lowest point of the

delta-plain crest is only -12 m. Colorado River gravels in the delta-plain sequence exposed in the Sierra de los Cucapas (Barnard

1966) suggest a significant hydrologic change in the Colorado River

during the late Pliocene or Pleistocene (Chapter 7), which may

represent a response to climatic change.

Lacustrine basin

Development of a closed lacustrine basin in the Salton Trough

is clearly a consequence of the Colorado delta, yet the change from

marine to lacustrine environments is not directly expressed in any

known stratigraphic section, and the oldest known lacustrine strata

(Borrego Formation) are not well dated. Arnal (1961) interpreted the

brackish-water microfaune in the Borrego Formation and fresh-water

microfaune in the Brawley Formation in terms of gradual freshening of

the Salton basin from marine to nonmarine conditions. However, such

long-term changes have not been clearly differentiated from short-term

fluctuations of salinity resulting from alternate filling and desic-

cation of the basin by switching of Colorado River distributaries.

Nowhere is the Borrego Formation known to directly overlie the Imperial

Formation, as implied by Arnal's (1961) scenario. Further clarification

of the lacustrine history of the Salton Trough will require better mapping, description, correlation, dating (as by magnetostratigraphy

and tephrochronology), and paleontologic study of lacustrine units. 408 Paleoclimates

A few lines of circumstantial evidence suggest that climate may have been somewhat more temperate during the Neogene and/or Pleistocene than at present. White and Downs (1961) interpreted the presence of Geomys, a grassland rodent now characteristic of the Great Plains, in the Hueso member of the FCV section as indicative of a less arid

climate. The presence of laurel as the dominant type of wood in delta- plain deposits (Diablo member) of the FCV section (Chapter 5) suggests

a broader distribution during the mid-Pliocene for this tree, which is not now common in the Colorado River drainage basin, and may indicate a more temperate climate (Remeika et al. in press). Colorado River gravel in late Pliocene/Pleistocene (?) delta-plain deposits in the Cucapas may indicate greater flood discharges in the Colorado River at that time (Barnard 1968). Evaporites were not observed in tidal flat depos-

its of the FCV deltaic sequence in the present study, whereas Thompson

(1968) described evaporites as characteristic of modern tidal flats on the Colorado delta. However, Miocene evaporites do occur in the FCV section (Fish Creek gypsum).

A mechanism for major climatic changes is not readily apparent. The arid regime that prevails today is largely a consequence of latitude and the rain-shadow effect of the Peninsular Ranges, a west- tilted basement block (Sylvester and Bonkowski 1979). If these ranges represent a rift-shoulder uplift related to proto-Gulf extension (Karig and Jensky 1973), the rain-shadow effect should date back at least to the middle Miocene. Higher sea-surface temperatures in the eastern 409 Pacific during the Neogene could conceivably have affected climate in the Salton Trough, although this effect is difficult to predict (Owen Davis, 1986, oral communication).

Stratigraphic implications of this study

Recognition of Colorado River sediment in the Neogene strati-

graphic record and consideration of sedimentological evidence were of

considerable benefit in clarifying chronology and paleoenvironments of

the Imperial and Palm Spring Formations. Therefore, it is recommended

that petrographic and sedimentologic observations be routinely included

in future stratigraphic studies of the Salton Trough and northern Gulf

of California. As a result of the present study, a locally-derived,

"pre-deltaic" phase of the Imperial Formation with time-transgressive

upper and lower contacts is established. Radiometric ages indicate that

the oldest locally-derived marine clastics are pre-Pliocene (Figure

7.1), whereas in places they interfinger with early Pliocene deltaic

sediments. The diverse molluscan-coral-echinoid faunas of the Imperial

occur mostly in this pre-deltaic phase, but biostratigraphy has so far

been inadequate to differentiate late Miocene from early Pliocene

assemblages. The deltaic phase of the Imperial, which contains most of

the known planktonic foraminiferal assemblages (except at Whitewater

Canyon), is exclusively early Pliocene. Marine macrofaunal assemblages

of the deltaic sequence are misleading about paleoenvironments because

of extensive reworking; a much better picture can be obtained from sedimentologic and taphonomic studies supplemented by micropaleontolo-

Or. However, direct comparison of the stratigraphic record of the 410

Pliocene delta with marine environments of the modern delta is often difficult, and many features of the stratigraphic record are not directly predictable from actualistic models of the Gulf of California and Colorado delta.

The overall stratigraphic succession in the western Salton

Trough reflects the tectonic history of the basin. The lower part of

the section is largely controlled by its rift-basin setting, and in-

cludes fanglomerates, alkalic and tholeiitic basalts, and pre-marine

evaporites, followed by marine sandstones and conglomerates deposited

as fan deltas. This rift phase is thought to manifest the onset of

interaction between the Pacific and North American plates. In accor-

dance with Walther's Law, the succession of paleoenvironments recorded

in the volumetrically dominant, fine-grained portion of the basin fill

(Figure 7.1) strongly resembles the geographic sequence of modern

environments from southeast to northwest in the northern Gulf and

Salton Trough: nondeltaic marine, deltaic marine, nonmarine delta

plain, and lacustrine. To some extent, this succession resulted from

initiation and early progradation of the Colorado delta, but it also

resulted from tectonic transport of the western Salton Trough past the

Colorado delta plain, which is anchored to the North American plate by

the fixed entry point of the Colorado River (Figure 8.15). Finally, an

important stratigraphic manifestation of the present transform basin

setting is the extensive angular unconformity currently developing due

to transpressive uplift of the Salton Trough margins. APPENDIX I

AERIAL PHOTOGRAPHIC COVERAGE

The study area is covered by monochrome aerial photography at

1:20,000 (nominal scale) taken for the USDA Soil Conservation Service on April 26, 1953. Other coverage is available, but the USDA-SCS photos are the most economical and widely used. The following table lists all photos used in the present study. The center of each photo is approximately located in terms of topographic quadrangle, the Township-Range

system, and latitude and longitude. Township-Range locations are

slightly abbreviated; for example, "NW,NE,4 T14S R7E" means "NW 1/4 of

NE 1/4 of Section 4, Township 14 East, Range 7 South". The Township-

Range system is not used in the northwestern portion of the study area.

In this table, flight lines are ordered from west to east, and

photos within each flight line are ordered from north to south. Each

photo (contact print) is 9" (23 cm) square, for a nominal coverage at

1:20,000 of 15,000' or 2.8 miles (4.6 km). Photo spacing along each

flight line is 1.0-1.2 miles (1.6-1.9 km). Spacing of flight lines is

variable, from 1.0 mile (1.6 km) to 2.5 miles (4.0 km).

411 412 Ordering Information

These photos can be ordered from the following address: Aerial Photography Field Office U. S. Dept. of Agriculture - A.S.C.S. P. 0. Box 30010 Salt Lake City, Utah 84125

When ordering, the following symbols should be specified:

For San Diego County:

symbol AXN PMA-29-53, 1:20,000, B&W, 4/26/53 survey.

For Imperial County:

symbol ABN PMA-29-53, 1:20,000, B&W, 4/26/53 survey.

At the time of the study, the price was $2.50/contact print, but prices

are subject to change. 413 SAN DIEGO COUNTY

Photo ID 7.5' Quadrangle Township-Range Lat. Long. AXN-16M-81 Whale Peak 33 ft 02.7' 116 15.9' AXN-16M-82 33 01.8' ff 116 15.9' AXN-16M-83 33 01.0' 11 116 15.9' AXN-16M-84 33 00.2' 116 15.9' AXN-16M-85 Agua Caliente Spgs NW,NE,4 T148 117E 32 59.3' 116 15.8' AXN-16M-86 NW,NE,9 T148 117E 32 58.5' 116 ft 16.0' AXN-16M-87 SW,SE,9 T148 117E 32 57.7' 116 16.0' ft AXN-16M-88 SW,SE,16 T148 117E 32 56.9' 116 16.0' t, AXN-16M-89 SW,SE,21 T148 117E 32 56.0' 116 16.0' AXN-16M-90 tt NE,NW,33 T148 117E 32 55.0' 116 16.1' AXN-16M-91 ff NW,NW,4 T158 R7E 32 54.1' 116 16.2' AXN-16M-92 11 NE,NW,9 T158 117E 32 53.1' 116 16.1' AXN-16M-135 Harper Canyon 33 02.0' 116 14.2' AXN-16M-134 33 01.0' 116 14.2' AXN-16M-133 Arroyo Tapiado 33 00.0' 116 14.2' AXN-16M-132 SW,NW,2 T148 117E 32 59.1' 116 14.3' AXN-16M-131 tt SW,NW,11 T148 117E 32 58.2' 116 14.4' AXN-16M-130 tt SE,NE,15 T148 117E 32 57.4' 116 14.6' AXN-16M-129 ft NE,SE,22 T148 117E 32 56.4' 116 14.6' AXN-16M-128 tt NE,SE,27 T148 117E 32 55.5' 116 14.7' AXN-16M-127 SW,NE,34 T148 117E 32 54.6' 116 14.8' AXN-16M-126 SE,NE,3 T158 117E 32 53.8' 116 14.6' AXN-16M-125 NE,SE,10 T158 117E 32 52.8' 116 14.7' AXN-16M-124 Sweeney Pass NE,SE,15 T158 117E 32 51.9' 116 14.8' AXN-16M-123 SW,NE,22 T158 117E 32 51.2' 116 14.8'

AXN-16M-195 Harper Canyon 33 03.2' 116 12.7' AXN-16M-196 1, 33 02.1' 116 12.8' AXN -16M-197 ft 33 01.2' 116 12.9' AXN-16M -198 ” 33 00.2' 116 12.9' AXN-16M-199 Arroyo Tapiado SE,NW,1 T148 117E 32 59.2' 116 13.0' AXN-16M-200 ft SE,NW,12 T148 117E 32 58.2' 116 13.1' AXN-16M-201 tt NE,SW,13 T148 117E 32 57.3' 116 13.0' AXN-16M-202 SW,SE,24 T148 117E 32 56.2' 116 12.9' AXN-16M-203 SW,SE,25 T148 117E 32 55.2' 116 12.9' AXN-16M-204 f t NW,NW,1 T158 117E 32 54.1' 116 12.9' AXN-16M-205 NE,NW,12 T158 117E 32 53.2' 116 13.0' AXN-16M-206 Sweeney Pass SE,NW,13 T158 117E 32 52.2' 116 13.0' AXN-16M-207 t, SE,NW,24 T158 117E 32 51.2' 116 13.0' AXN-16M-208 tt NW,SW,25 T158 117E 32 50.2' 116 13.0' AXN-16M-209 tt NW,SW,36 T158 117E 32 49.2' 116 13.1'

AXN-4M-210 Harper Canyon 33 02.8' 116 10.4' AXN-4M-209 11 33 01.9' 116 10.4' AXN-4M-208 It 33 01.1' 116 10.5' 414 Photo ID 7.5' Quadrangle Township-Range Lat. I2ag AXN-4M-207 Arroyo Tapiado 32 AXN-4M-206 59.9' 116 10.5' NE,SE,5 T14S R8E 32 58.9' 116 AXN-4M-205 11 10.4' SE,SE,8 T148 R8E 32 57.9' 116 AXN-4M-204 tt 10.4' SE,SE,17 T14S R8E 32 56.9' AXN-4M-203 116 10.5' NE,NE,29 T14S R8E 32 55.8' 116 AXN-4M-202 ft 10.6' SW,NE,32 T14S R8E 32 54.8' 116 10.6' AXN-4M-201 If SW,NE,5 T15S R8E 32 53.7' 116 10.6' AXN-4M-200 It NW,SE,8 T15S R8E 32 52.7' 116 10.7' AXN-4M-199 Sweeney Pass SW,SE,17 T15S R8E 32 51.7' 116 AXN-4M-198 10.7' NW,NE,29 T1SS R8E 32 50.6' 116 10.7' AXN-4M-197 tt SE,NE,32 T15S R8E 32 49.5' 116 10.6' AXN-4M-196 SW,NE,5 T16S R8E 32 48.6' 116 10.6'

AXN-15M-83 Harper Canyon 33 02.6' 116 09.1' AXN-15M-84 33 01.6' 116 09.1' tt AXN-15M-85 33 00.5' 116 09.1' AXN-15M-86 Arroyo Tapiado 32 59.5' 116 09.1' AXN-15M-87 NW,NW,10 T148 R8E 32 58.5' 116 09.1' AXN-15M-88 SW,NW,15 T14S R8E 32 57.4' 116 09.1' AXN-15M-89 tt NW,SW,22 T14S R8E 32 56.4' 116 09.1' AXN-15M-90 It NW,SW,27 T14S R8E 32 55.4' 116 09.2' AXN-15M-91 tt SW,SW,34 T14S R8E 32 54.4' 116 09.2' AXN-15M-92 tf NW,NW,10 T15S R8E 32 53.3' 116 09.3' AXN-15M-93 Sweeney Pass NE,NE,16 T15S R8E 32 52.3' 116 09.4' AXN-15M-94 t, NE,NE,21 T15S R8E 32 51.3' 116 09.4' AN-15M-95 SE,NE,28 T15S R8E 32 50.3' 116 09.4' AXN-15M-96 tt NE,SE,33 T15S R8E 32 49.2' 116 09.4'

AXN-15M-136 Borrego Mtn SE NW,NE,34 T12S R8E 33 05.2' 116 07.2' AXN-15M-135 SW,NW,1 T13S R8E 33 04.2' 116 07.2' AXN-15M-134 NE,SE,11 T13S R8E 33 03.2' 116 07.2' AXN-15M-133 ff SW,SW,13 T13S R8E 33 02.3' 116 07.1' AXN-15M-132 tt SW,SW,24 T13S R8E 33 01.3' 116 07.1' AXN-15M-131 ft SW,SW,25 T13S R8E 33 00.3' 116 07.1' AXN-15M-130 Carrizo Mtn NE NE,NE,2 T14S R8E 32 59.4' 116 07.2' AXN-15M-129 tt NE,NE,11 T14S R8E 32 58.4' 116 07.2' AXN-15M-128 NE,NE,14 T14S R8E 32 57.5' 116 07.2' AXN-15M-127 11 SE,NE,23 T14S R8E 32 56.5' 116 07.2' AXN-15M-126 SE,NE,26 T14S R8E 32 55.6' 116 07.2' AXN-15M-125 tt NE,SE,35 T14S R8E 32 54.6' 116 07.3' AB-15M-124 It NE,SE,2 T15S R8E 32 53.7' 116 07.3' AXN-15M-123 tf NE,SE,11 T15S R8E 32 52.6' 116 07.4' AXN-15M-122 Carrizo Mtn SE,SE,14 T15S R8E 32 51.6' 116 07.4' AXN-15M-121 11 NE,NE,26 T15S R8E 32 50.6' 116 07.4' AXN-15M-120 ft NE,NE,35 T15S R8E 32 49.7' 116 07.4' AXN-15M-119 ts NE,NE,2 T16S R8E 32 48.8' 116 07.4' 415 IMPERIAL COUNTY

Photo ID 7.5' Quadrangle Township-Range Lat. Long.

ABN-13M-124 Borrego Mtn SE SW,SW,36 T12S,R8E 33 04.5' 116 05.8' ABN-13M-123 SW,NW,7 T13S,R9E tt 33 03.6' 116 05.9' ABN -13M -122 SW,NW,18 T13S,R9E 33 02.6' 116 05.9' ABN-13M -121 NW,SW,19 T13S,R9E tt 33 01.6' 116 05.9' ABN -13M -120 NE,SW,30 T13S,R9E 33 00.7' 116 05.8' ABN-13M-119 Carrizo Mtn NE tt NE,SW,31 T13S,R9E 32 59.7' 116 05.8' ABN-13M-118 SE,SW,6 T14S,R9E 32 58.7' 116 05.8'

ABN-13M -115 ft NW,NW,6 T14S,R9E 32 59.4' 116 06.0' ABN -13M -114 ft NW,NW,7 T14S,R9E 32 58.4' 116 06.1' ABN -13M-113 tt NW,NW,18 T14S,R9E 32 57.5' 116 06.1' ft ABN -13M -112 NW,NW,19 T14S,R9E 32 56.6' 116 06.1' ABN-13M -111 't NW,NW,30 T14S,R9E 32 55.9' 116 06.1' tt ABN -13M -110 NW,NW,36 T14S,R8E 32 54.8' 116 06.2' ABN-13M-109 ft SE,NE,1 T15S,R8E 32 53.9' 116 06.2' ABN -13M -108 tt SE,NE,12 T15S,R8E 32 53.0' 116 06.3' ABN-13M -107 Carrizo Mtn SE,NE,13 T15S,R8E 32 52.1' 116 06.3' ABN -13M -106 tt SE,NE,24 T15S,R8E 32 51.1' 116 06.2' ABN-13M-105 t NE,SE,25 T15S,R8E 32 50.2' 116 06.2' ABN -13M-104 NE,SE,36 T15S,R8E 32 49.2' 116 06.2' ABN-13M-103 NE,SE,1 T16S,R8E 32 48.3' 116 06.2'

ABN -12M -172 Borrego Mtn SE NE,NE,29 T13S R9E 33 01.1' 116 04.2' tt ABN -12M -173 NE,NE,32 T13S R9E 33 00.2' 116 04.2' ABN -12M -174 Carrizo Mtn NE NE,NE,5 T14S R9E 32 59.3' 116 04.1' ABN-12M -175 t, NW,NW,9 T14S R9E 32 58.4' 116 03.9' ABN -12M -176 11 NE,NE,17 T14S R9E 32 57.5' 116 04.1' ABN-12M-177 NE,NE,20 T14S R9E 32 56.6' 116 04.1' ABN -12M -178 tt NE,NE,29 T14S R9E 32 55.7' 116 04.2' ABN -12M-179 f SE,NE,32 T14S R9E 32 54.8' 116 04.2' ABN -12M -180 /t NW,NW,5 T15S R9E 32 54.0' 116 04.2' ABN -12M -181 tf NW,NW,8 T15S R9E 32 53.1' 116 04.2' ABN -12M-182 Carrizo Mtn NE,NE,18 T15S R9E 32 52.2' 116 04.3' ABN -12M -183 t t NE,NE,19 T15S R9E 32 51.2' 116 04.3' ABN-12M -184 f? SE,NE,30 T15S R9E 32 50.2' 116 04.3' ABN -12M -185 tt SE,NE,31 T15S R9E 32 49.4' 116 04.3' ABN-12M -186 NE,NE,6 T16S R9E 32 48.5' 116 04.5' ABN -12M -187 It SE,NE,7 T16S R9E 32 47.6' 116 04.5' 416 Photo ID 7.5' Quadrangle Township-Range Lat. Long. ABN-12M-204 Borrego Mtn SE NE,NE,27 T13S R9E 33 01.0' 116 02.0' ABN-12M-203 Carrizo Mtn NE NW,SE,35 T13S R9E 32 59.7' ABN-12M-202 116 01.8' NE,SE,3 T14S R9E 32 58.9' 116 01.9' ABN-12M-201 It NW,SW,11 T14S R9E 32 57.9' 116 01.9' ABN-12M-200 11 SW,SW,14 T14S R9E 32 56.9' 116 01.8' ABN-12M-199 ft NW,NW,26 T14S R9E 32 55.9' 116 01.8' ABN-12M-198 NW,NW,35 T14S R9E 32 54.8' 116 01.7' ABN-12M-105 ft NE,SE,27 T14S R9E 32 55.5' 116 ft 02.1' ABN-12M-104 NE,SE,34 T14S R9E 32 54.5' 116 02.1' ABN-12M-103 11 NW,SW,3 T15S R9E 32 53.7' 116 02.1' ABN-12M-102 11 SW,NW,10 T15S R9E 32 52.7' 116 02.1' ABN-12M-101 Carrizo Mtn SE,NE,16 T15S R9E 32 51.8' 116 02.2' ABN-12M-100 NE,SE,21 T15S R9E 32 50.9' 116 02.5' ABN-12M-99 11 NE,SE,28 T15S R9E 32 50.0' 116 02.4' ABN-12M-98 ft NW,SE,33 T15S R9E 32 49.2' 116 02.5' ABN-12M-97 tt NW,SE,4 T16S R9E 32 48.3' 116 02.5' ABN-12M-96 ft NW,SE,9 T16S 119E 32 47.4' 116 02.6' ABN-12M-95 Pt NW,SE,16 T16S 119E 32 46.6' 116 02.6'

ABN -12M-65 Carrizo Mtn NE NW,NE,1 T14S 119E 32 59.5' 116 00.3' ABN -12M -66 SW,SE,1 T14S 119E 32 58.7' 116 00.3' ABN -12M -67 ft SW,SE,12 T14S 119E 32 57.9' 116 00.3' ABN-12M -68 SW,SE,13 T14S 119E 32 57.1' 116 00.3' ABN-12M -69 SW,SE,24 T14S 119E 32 56.2' 116 00.3' ABN -12M -70 tt SW,SE,25 T14S 119E 32 55.3' 116 00.3' ABN-12M -71 11 SW,SE,36 T14S 119E 32 54.3' 116 00.3' ABN-12M -72 ft NE,SE,2 T15S 119E 32 53.5' 116 00.3' ABN-12M -73 tt NW,SE,11 T15S 119E 32 52.5' 116 00.3' ABN -12M-74 ft SE,SE,14 T15S 119E 32 51.6' 116 00.3' ABN-12M-75 ft NE,SE,23 T15S 119E 32 50.8' 116 00.3' ABN-12M -76 If NW,SE,26 T15S R9E 32 49.9' 116 00.4' ABN-12M-77 tt NW,SE,35 T15S 119E 32 49.0' 116 00.5' ABN -12M -78 ft NW,SE,2 T16S R9E 32 48.2' 116 00.5' ABN-12M -79 ft NW,SE,11 T16S 119E 32 47.3' 116 00.4' ABN-12M -80 ft SE,SE,14 T16S 119E 32 46.4' 116 00.3' 417 Photo ID 7.5, Quadrangle Township-Range Lat. Long.

ABN-10M-198 Plaster City NW NE,NW,31 T14S R1OE 32 54.7' 115 59.1' ABN-10M-199 SE,NE,1 T15S R9E 32 53.7' 115 59.2' ABN-10M-200 SE,NE,12 T15S R9E 32 52.7' 115 59.1' ABN-10M-201 Painted Gorge NW,SW,18 T15S 1110E 32 51.7' 115 59.0' ABN-10M-202 NW,SW,19 T15S 1110E 32 50.8' 115 59.0' ABN-10M-203 NW,SW,30 T15S R1OE 32 49.9' 115 58.9' ABN-10M-204 SW,SW,31 T15S R1OE tt 32 49.0' 115 58.9' ABN-10M-205 SW,SW,6 T16S R1OE 32 48.0' 115 58.9' ABN-10M-206 NW,NW,18 T16S R1OE 32 47.0' 115 58.8'

ABN -12M -4 Plaster City NW SW,SW,32 T13S R1OE 32 59.3' 115 57.9' ABN-12M -3 tt NW,NW,8 T14S R1OE 32 58.3' 115 57.9'

ABN-11M-231 ft SE,SE,6 T14S R1OE 32 58.4' 115 58.3' ABN-11M-230 tt SE,SE,7 T14S R1OE 32 57.5' 115 58.2' ABN-11M-229 It NE,NE,19 T14S R1OE 32 56.6' 115 58.2' ABN-11M-228 11 NE,NE,30 T14S 1110E 32 55.6' 115 58.2' ABN-11M-227 11 NE,NE,31 T14S R1OE 32 54.7' 115 58.1' ABN-11M-226 tt NW,NW,5 T15S 1110E 32 53.8' 115 57.9' APPENDIX II

PALEOCURRENT DATA

Paleocurrent data are presented as azimuths after rotation to correct for present structural tilt. True paleocurrent directions are probably different, due to inferred -35° of postdepositional clockwise rotation (Chapter 8; Johnson et al. 1983). Data were obtained from the following paleocurrent indicators:

FC - flute casts

GC - groove casts

SM - other bidirectional sole marks

XS - foresets in cross-stratification

RL - foresets in small-scale ripple lamination

RM - small-scale ripple marks on bedding plane

PL - parting lineation (if unidirectional, sense inferred

from adjacent ripple lamination)

U - unidirectional

B - bidirectional

Data are subdivided by stratigraphic unit and in some cases by

geographic area and/or provenance as well.

418 419 LYCIUM MEMBER (LOWER TURBIDITE UNIT)

Location Indicator Azimuth Location Indicator Azimuth

WC-m-16 GC B 177 CC-m-1 SM B 166 tt SM B 170 te SM B 155 et FC U 194 et SM B 100 ff SM B 196 et SM B 148 WC-m-17 SM B 172 te RL U 212

LYCIUM MEMBER (UPPER TURBID1TE UNIT)

Location Indicator Azimuth Location Indicator Azimuth

0S-m-5 FC U 055 OS-m-21 FC U 105 te FC U 080 LW-1-19 FC U 060 11 FC U 045 LW-1-22 FC U 075 tt FC U 070 OS-2-5 FC U 055 FC U 075 11 GCB 085 tt FC U 075 LC-m-? RM U 053 te FC U 065 SM-2-28 GCB 150 et FC U 075 SM-2-30-33 GCB 072 tt FC U 070 tf GCB 070 11 RM B 070 11 GCB 035 FC U 082 SM-2-38 GCB 074 ft FC U 065 SM-2a-4 GCB 085 OS-m-18 FC U 095 SM-3-10-15 GC B 070 SM B 115 ft GC B 075 tt SM B 100 SM-3-40 GCB 130 ft FC U 120 SM-3-42 RM B 025 tf GCB 135 SM-3-48 RM U 025 ft FC U 090 SM-3-58 GC B 137 OS-m-19 SM B 110 SM-3-64 FC U 120 ff XS U 085

STONE WASH MEMBER (DISTAL FACIES)

Location Indicator Azimuth Location Indicator Azimuth

LW-1-104 FC U 090 OS-1-21 FC U 100 0S-1-19 FC U 085 OS-1-275 FC U 085

KAMIA CREEK MEMBER

Location Indicator Azimuth Location Indicator Azimuth

FC B 060 FC B 065 420 WIND CAVES MEMBER (OYSTER SHELL WASH)

Location Indicator Azimuth Location Indicator Azimuth

0S-2-45-48 PC U 120 OS-m-11 PC U 205 GC B 106 PC U 222 SM B 115 GC B 190 OS-1-72.5 SM B 205 PC U 190 0S-m-1 PC U 190 PC U 185 0S-m-2 FC U 110 GC B 198 PC U 110 GC B 218 PC U 120 OS-m-12 SM B 260 PC U 120 OS-m-13 GC B 248 PC U 135 FC U 215 0S-m-3 PC U 130 PC U 228 0S-m-6 GC B 204 FC U 220 GC B 232 PC U 222 GC B 095 GC B 222 0S-m-7 GC B 165 OS-m-14 GC B 235 SM B 165 GC B 190 0S-m-8 PC U 190 OS-m-15 FC U 195 0S-m-9 SM B 115 GC B 210 OS-m-10 GC B 180 OS-m-16 GC B 205 GC B 080 OS-m-17 SM B 220 GC B 182 PC U 218 OS-m-11 PC U 202 SM B 210 FC U 210

WIND CAVES MEMBER (NORTH FORK)

Location Indicator Azimuth Location Indicator Azimuth

NP-1-2 SM B 165 NF-1-143.5 RM U 095 NF-1-38 PC U 260 NF-1-177 PC U 195 NP-1-42 PC U 251 NF-1-183 PC U 205 NF-1-100 RM U 200 NF-1-184 PC U 220 NF-1-101 PL U 218 NF-1-195 GC B 210

WIND CAVES MEMBER (HEAD OF SPLIT MOUNTAIN GORGE)

Location Indicator Azimuth Location Indicator Azimuth

FC-5-29 PC U 200 PC-5-60 RL U 175 FC-5-30 PC U 185 SM-1-55.5 FC U 140 FC-5-47 GC B 195 SM-1-82 PC U 175 FC-5-51 GC B 180 SM-1-93 PC U 185 FC-5-52 PC U 200 5M-1-84 RM U 205 421 WIND CAVES MEMBER (WIND CAVES WASH)

Location Indicator Azimuth Location Indicator Azimuth

WC-m-1 FC U 188 WC-m-9 FC U 205 ft FC U 178 't FC U 200 FC U 215 WC-m-10 FC U 160 GCB 190 11 SM B 205 FC U 175 SM B 185 WC-m-2 FC U 175 WC-m-11 FC U 225 FC U 155 11 FC U 210 SM B 204 WC-m-12 FC U 140 SM B 200 11 FC U 162 11 SM B 195 tt FC U 162 WC-m-3 FC U 182 f FC U 165 WC-m-4 FC U 192 11 FC U 165 FC U 180 FC U 160 WC-m-5 FC U 170 11 FC U 160 FC U 170 ft FC U 182 SM B 140 ft FC U 160 WC-m-6 FC U 150 ft FC U 168 WC-m-7 FC U 175 WC-m-13 SM B 180 11 SM B 160 ft SM B 190 FC U 152 ft SM B 162 FC U 210 ft GCB 175 FC U 182 WC-m-14 FC U 160 FC U 185 ft FC U 182 WC-m-8 FC U 200 ft FC U 192 FC U 250 ft SM B 165 11 FC U 192 WC-m-15 FC U 140 WC-m-9 FC U 202 ft FC U 125

WIND CAVES MEMBER (CAIRN WASH)

Location Indicator Azimuth Location Indicator Azimuth

ER-1-17 FC U 132 EK-1-82 GC B 155 ER-1-18 FC U 147 EK-1-90 FC U 158 ER-1-30 GC B 148 EX-1-91 168 ft FC U GC B 150 EK-1-100 GC B 130 ER-1-34 GC B 155 EX-1-180 FC U 160 EK-1-36 GC B 170 EX-m-8 FC U 140 ER-1-37 GC B 140 ,, FC U 155 ER-1-44 GC B 132 ,, FC U 120 ER-1-59 SM B 170 EX-m-9 SM B 200 422 LAVENDER CANYON MEMBER

Location Indicator Azimuth Location Indicator Azimuth

Lavender Cyn FC U 125 Lavender Cyn FC U 160 FC U 150 RR-1-72 FC U 175 tt FC U 142 RR-1-80 FC U 135 ff FC U 210 RR-1-89 PL B 130

DEGUYNOS MEMBER (ELEPHANT KNEES AREA)

Location Indicator Azimuth Location Indicator Azimuth

EK-2-5-10 XS U 164 ETC-m-1 XS U 248 tt XS U 161 ETC-m-2 XS U 178 XS U 169 it XS U 215 EK-2-30-40 XS U 195 tt XS U 215 t, XS U 180 tt XS U 212 XS U 175 XS U 210 XS U 251 ff XS U 004 It XS U 183 t, XS U 344 1 1 XS U 199 EK-m-3 XS U 187 XS U 187 tt XS U 171 t t XS U 180 t, XS U 358 XS U 197 tt XS U 348 tt XS U 254 tt XS U 315 1 1 XS U 247 ETC-m-4 XS U 183 EK-2 -43 -47 XS U 181 tt XS U 163 XS U 166 tf XS U 180 11 XS U 155 XS U 173 It XS U 155 it XS U 129 EK-2-44-47 XS U 171 tt XS U 127 t, XS U 175 ft XS U 137 t t XS U 186 tt XS U 164 ft XS U 180 tt XS U 154 EK-1-130-133 XS U 243 tt XS U 121 XS U 224 ETC-m--5 XS U 174 tt XS U 228 XS U 195 XS U 228 t, XS U 175 ETC-in-1 XS U 216 ETC-m-6 XS U 116 XS U 272 ti XS U 113 1 1 XS U 231 tt XS U 116 It XS U 265 ETC-m-7 XS U 225 ft XS U 287 it XS U 177 423 DEGUYNOS MEMBER (NORTH FORK) Location Indicator Azimuth Location Indicator Azimuth

NF-3-15-17 XS U 221 NF-4-153-154 PM U 236 XS U 161 RM U 233 XS U 222 NF-5-55 XS U 172 XS U 206 NF-5-59-61 XS U 204 XS U 191 XS U 236 XS U 244 NF-5-86-88 XS U 080 XS U 206 XS U 074 NF-3-71-73 XS U 272 I/ XS U 009 XS U 191 XS U 049 XS U 158 XS U 342 XS U 172 XS U 158 NF-3-86-91 XS U 145 XS U 038 XS U 167 XS U 168 XS U 018 XS U 049 XS U 016 XS U 077 XS U 144 XS U 011 XS U 161 XS U 067 XS U 134 XS U 067 XS U 114 XS U 098 NF-3-133-135 XS U 208 XS U 112 XS U 171 XS U 063 XS U 230 NF-5-109-110 PM U 227 XS U 185 NF-5-114-115 XS U 182 NF-4-45-49 XS U 178 XS U 237 NF-4-153-154 PM U 238 XS U 196 PM U 244 XS U 162 RM U 255

DEGUYNOS MEMBER (CORAL WASH)

Location Indicator Azimuth Location Indicator Azimuth

XS U 009 XS U 048 XS U 009 XS U 015 XS U 356

424 DEGUYNOS MEMBER (FISH CREEK WASH)

Location Indicator Azimuth Location Indicator Azimuth

FC-6-230-232 XS U 175 FC-1-25-28 XS U 239 tt If XS U 194 XS U 178 tt XS U 199 XS U 234 /I tt XS U 355 XS U 188 H XS U 160 XS U 155 FC-7-109-112 XS U 185 It XS U 179 ft It XS U 082 XS U 323 rt XS U 231 FC-1-35 XS U 280 ,t XS U 131 FC-1-43-46 XS U 315 tt XS U 099 It XS U 218 it ,, XS U 273 XS U 343 u ,, XS U 207 XS U 258 ft XS U 123 FC-1-84 XS U 202 FC-1-205-207 XS U 321

DEGUYNOS MEMBER (RED ROCK CANYON)

Location Indicator Azimuth Location Indicator Azimuth

RR-1-152-155 XS U 155 RR-1-152-155 XS U 112 XS U 159 'I U 175 r, XS U 136 425 CAMELS HEAD MEMBER (FISH CREEK WASH)

Location Indicator Azimuth Location Indicator Azimuth FC-1-264-266 XS U 202 FC-2-113-114 tt XS U 212 XS U 191 FC-3-0-5 XS U 177 tt XS U 194 XS U 184 FC-2-73-76 XS U 178 tt XS U 281 XS U 154 XS U 223 XS U 128 tt XS U 244 XS U 148 XS tt U 229 XS U 206 FC-3-16-18 XS U 278 FC-2-101-102 XS U 187 XS U 252 XS U 232 XS U 223 FC-2-108-112 XS U 298 FC-3-33-35 XS U 099 XS U 091 XS U 132 XS U 234 XS U 033 XS U 285 FC-3-64 XS U 082 XS U 224 FC-3-94 XS U 000 XS U 279 FC-3-102-104 XS U 152 XS U 224 ff XS U 347 tt XS U 211 XS U 065 FC-2-113-114 XS U 231 FC-3-132 XS U 044 tt XS U 321 FC-3-142 XS U 185 tt XS U 217

CAMELS HEAD MEMBER (MOLLUSK WASH)

Location Indicator Azimuth Location Indicator Azimuth

MW-1-45-46 XS U 224 MW-1-111-115 XS U 246 XS U 212 XS U 265 MW-1-65-66 XS U 068 XS U 271 XS U 068 XS U 156 MW-1-76-80 XS U 252 XS U 015 XS U 254 XS U 198 ft XS U 259 XS U 222 XS U 302 MW-1-117-119 RL U 011 XS U 321 RL U 333 MW-1-107-108 XS U 078 t, RL U 197 426 CAMELS HEAD MEMBER (NF-6, NORTH FORK)

Location Indicator Azimuth Location Indicator Azimuth

NF-6-5-15 RL U 170 NF-6-31-33 XS U 191 PLU 194 t, XS U 208 /I PLU 135 tt XS U 186 tt RL U 144 NF-6-31-33 XS U 184 tt PLU 159 NF-6-44 RL U 203 tt PLU 187 If RL U 194 PLU 179 tt U 204 PLU 183 PLU 344 tt PLU 176 t, PLU 000 PLU 036 11 PLU 334 t, PLU 182 NF-6-45-48 XS U 202 ft RL U 174 XS U 200 t, PLU 026 XS U 203 tt PLU 166 t, XS U 191 PLU 154 NF-6-49-50 XS U 179 PLU 262 If XS U 174 tt PLU 206 XS U 187 PLU 195 XS U 206 PLU 237 NF-6-58-59 PM U 335 PL U 064 PMU 175 PLU 189 NF-6-80-82 PM U 210 PLU 128 tt RM U 175 PLU 198 NF-6-94-97 XS U 199 tl PLU 061 XS U 207 t, RL U 046 XS U 217 PLU 345 XS U 206 t, PLU 020 XS U 194 t, PLU 353 XS U 199 NF-6-19 XS U 002 XS U 184 NF-6-25 XS U 268 t, XS U 203 XS U 211 XS U 208 XS U 250 PLU 178 XS U 183 PLU 169 ff XS U 215 PLU 201 NF-6-27-30 XS U 192 tf PLU 222 XS U 175 PLU 182 XS U 192 PL B 201 XS U 139 PL B 908 NF-6-31-33 XS U 219 PL B 205 t, XS U 198 NF-6-97-99 XS U 154 XS U 191 XS U 169 tt XS U 195 XS U 195 t, XS U 209 XS U 171 t, XS U 189 XS U 215 XS U 21] XS U 169 XS U 208 XS U 200 427 CAMELS HEAD MEMBER (NF-6, NORTH FORK) Location Indicator Azimuth Location Indicator Azimuth

NF -6 -97 -99 XS U 188 NF-6-97-99 XS U 220 XS U 184 tf XS U 196 XS U 172 ft XS U 180

CAMELS HEAD MEMBER (NF-7, NORTH FORK)

Location Indicator Azimuth Location Indicator Azimuth

NF-7-5 PM U 207 NF-7-90-92 PM U 180 RM U 160 RMU 188 NF-7-13 PM U 195 RM U 189 RM U 192 NF-7-92-94 XS U 181 NT-7-16 XS U 138 tt XS U 171 XS U 162 tt XS U 157 NF-7-83 XS U 182 XS U 188 XS U 118 tt XS U 204 ff XS U 212 XS U 231 ft RM B 200 tt XS U 191 RM U 189 XS U 204 NF-7-87-90 XS U 176 XS U 205 XS U 129 XS U 155 XS U 153 XS U 129 XS U 192 tt XS U 193 Ft XS U 201 NF-7-98-99 XS U 129 XS U 169 11 XS U 158 XS U 187 XS P 172 XS U 168 NF-7-10 4 PM U 200 XS U 192 NF-7-108-109 XS U 163 XS U 166 tt XS U 245 XS U 163 XS U 211 XS U 168 XS U 215 NF-7-90-92 RN U 180 11 XS U 132 ft RM U 185 NF-7-133-134 XS U 120 RM U 185 RM U 205

CAMELS HEAD MEMBER (JACKSON FORK)

Location Indicator Azimuth Location Indicator Azimuth

RM U 020 CT-2-183-184 XS U 199 RM U 030 CT-2-220 XS U 187 RM U 035 XS U 213 CT-2-130-131 XS U 256 XS U 208 ft XS U 296 XS U 204 ft XS U 189

428 DIABLO MEMBER (MOLLUSK WASH)

Location Indicator Azimuth Location Indicator Azimuth

MW-1-130-133 XS U 308 XS U 138 tt XS U 260 XS U 175 MW-1-138 XS U 237 XS U 042 MW-1-136 RL U 002 XS U 093 RL U 009 MW-2-7-9 XS U 216 MW-1-154-159 XS U 248 XS U 200 XS U 213

DIABLO MEMBER (NORTH FORK)

Location Indicator Azimuth Location Indicator Azimuth

NF-7-153-154 XS U 208 NF-7-162-164 XS U 004

ft XS U 162 XS U 066

11 XS U 188 NF-7-176-178 XS U 142

NF-7-157-159 XS U 215 tt XS U 147

XS U 225 tt XS U 154

tt XS U 172 Tt XS U 136

XS U 194 It XS U 176 XS U 215 NF-7-197 XS U 172 XS U 211 429 DIABLO MEMBER (FISH CREEK WASH)

Location Indicator Azimuth Location Indicator Azimuth

FC-3-176 RL U 320 FC-4-134-139 RL U 310 FC-3-203 PL B 140 RL U 345 FC-3-205-206 XS U 071 FC-4-141-142 PL B 090 t, XS U 085 PL B 085 FC-3-231-236 XS U 171 FC-4a-2-4 XS U 084 XS U 272 t, XS U 211 tt XS U 219 XS U 141 t, XS U 216 FC-4a-5-8 PL B 108 t, tt XS U 198 PL B 155 t, RL U 340 PL B 125 FC-3-283-287 XS U 193 PL B 142 tt It XS U 207 PL B 143 XS U 191 tt PL B 122 XS U 231 XS U 171 FC-3a-0-3 XS U 237 FC-4a-8-9 XS U 143 XS U 289 tt XS U 165 FC-3-102-104 XS U 346 FC-4a-15 RL U 132 t, XS U 067 RL U 135 FC-3-94 XS U 358 FC-4a-18 XS U 049 FC-3-132 XS U 044 FC-4a-43 PL B 160 FC-3-142 XS U 184 FC-4a-55 PL B 185 FC-3-174 RL U 320 FC-4a-63-66 XS U 204 FC-3-229 RL U 340 XS U 198 FC-3-249 XS U 265 XS U 182 FC-3-260-261 XS U 058 FC-4a-80-81 XS U 230 t, XS U 049 tt XS U 216 FC-4-40 RM U 235 tt XS U 182 FC-4-72-73 XS U 282 FC-4a-87-90 PL B 070 XS U 189 tt XS U 044 XS U 165 t, XS U 349 FC-4-102 XS U 291 tt RL U 080 FC-4-128-131 PL B 230 FC-4b-0-2 XS U 051 tt PL B 225 tt XS U 059 tt PL B 230 FC-4b-12-14 XS U 218 PL B 218 tt XS U 228 tt PL B 240 FC-4b-18 PL B 175 PL B 232 FC-4b-58-64 XS U 082 FC-4-139 PL B 322 XS U 184 FC-4-134-139 RL U 000 tt PL B 180 RL U 345 FC-4b-32 PL B 214 RL U 338

430 DIABLO MEMBER (BLACKWOOD BASIN WASH)

Location Indicator Azimuth Location Indicator Azimuth

BW-m-1 (PC1) XS U 054 BW-m-2 (PC1) XS U 196 ft XS U 030 XS U 123 PL B 125 XS U 188 BW-m-1 (PC2) XS U 057 BW-m-2 (PC2) XS U 176 XS U 077 tf XS U 165 XS U 347 XS U 178 XS U 156 XS U 144 XS U 183 BW-m-3 XS U 149 XS U 158 XS 166 XS U 169 XS U 134 XS U 055 XS U 186 BW-m-1 (PC3) XS U 166 XS U 204 XS U 182 BW-m-4 XS U 177 BW-m-1 (PC4) XS U 113 XS U 182 XS U 115 11 XS U 138 XS U 107 XS U 139 BW-m-2 (PC1) XS U 178

DIABLO MEMBER (ARROYO TAPIADO)

Location Indicator Azimuth Location Indicator Azimuth

AT-m-9 XS U 174 AT-m-10 XS U 131 If XS U 228 tf XS U 141 tt XS U 198 tt XS U 131 ft XS U 194 If XS U 160 ft XS U 176 tt XS U 159 ft XS U 202 It XS U 143 tt XS U 139 11 XS U 156 tt XS U 180 tt XS U 106 ft XS U 147 ft XS U 075 XS U 175 XS U 116 XS U 168 11 XS U 182 XS U 174 11 XS U 164 XS U 157 tf XS U 158 XS U 142 it XS U 175 XS U 133 I XS U 193 XS U 108 ff XS U 215 XS U 149 ft XS U 188 XS U 149 tt XS U 220 XS U 165 11 XS U 149 tt XS U 109 431 DIABLO MEMBER (ARROYO TAPIADO) Location Indicator Azimuth Location Indicator Azimuth AT-m-11 XS U 159 AT-m-14 XS U 112 11 XS U 112 XS U 080 tt XS U 104 XS U 092 tt XS U 120 XS U 086 tt XS U 123 XS U 087 XS U 140 XS U if 070 XS U 140 )(SU 051 AT-m-12 (PC1) XS U 132 tf XS U 067 XS U 155 fl XS U 097 XS U 177 XS U 076 XS U 186 AT-m-15 XS U 145 XS U 147 tf XS U 115 XS U 169 XS U 133 tt XS U 120 XS U 105 It tt XS U 138 XS U 115 I f XS U et 122 XS U 141 XS U 144 AT-m-17 XS U 043 XS U 106 tt XS U 080 XS U 078 XS U 146 XS U 102 XS U 080 XS U 111 AT-m-18 XS U 078 XS U 098 XS U 088 tt XS U 116 tt XS U 240 AT-m-12 (PC2) XS U 179 XS U 240 XS U 217 AT-m-19 XS U 084 XS U 309 XS U 075 XS U 213 XS U 108 XS U 180 tt XS U 058 XS U 188 XS U 115 XS U 180 XS U 121 tt XS U 175 Ft XS U 098 f XS U 166 XS U 121 AT-m-13 XS U 045 tt XS U 109 XS U 027 AT-m-20 XS U 191 XS U 052 XS U 184 XS U 101 XS U 118 XS U 031 XS U 201 XS U 059 tf XS U 197 tt XS U 078 AT-m-21 XS U 045 XS U 063 XS U 086 It XS U 072 XS U 068 tf XS U 049 XS U 052 XS U 046 XS U 052 432 DIABLO MEMBER (DIABLO CANYON)

Location Indicator Azimuth Location Indicator Azimuth

DC-1-1-3 XS U 217 DC-4-9-11 et XS U 136 XS U 186 XS U 131 XS U 224 XS U 112 XS U 226 XS U 164 DC-1-10-12 XS U 096 DC-4-11-13 XS U 172 te XS U 088 XS U 133 t, XS U 061 DC-5-3-5 XS U 232 DC-1-12-14 XS U 159 XS U 121 te tt XS U 106 XS U 101 XS U 089 PL B 142 ttt XS U 135 DC-5-9-11 XS U 141 te tt XS U 191 XS U 149 tt te XS U 105 XS U 125 tt t, XS U 154 XS U 146 tt XS U 124 XS U 142 DC-1-18 XS U 111 DC-5-14 XS U 109 DC-1-23-25 XS U 128 XS U 176 XS U 126 XS U 135 XS U 145 DC-5-22-27 XS U 162 XS U 186 XS U 149 DC-2-0-4 XS U 194 et XS U 165 XS U 163 XS U 139 XS U 194 XS U 162 XS U 198 XS U 171 tt XS U 137 XS U 166 tt XS U 151 XS U 145 XS U 155 XS U 169 et XS U 165 XS U 169 XS U 185 XS U 134 tt XS U 185 XS U 166 te XS U 164 XS U 108 DC-2-17-19 XS U 229 XS U 131 tt PL B 250 ff XS U 149 tt PL B 240 ff XS U 102 PL B 222 tf XS U 120 DC-3-1-3 XS U 120 t, XS U 174 e, XS U 136 te XS U 160 tt XS U 151 tt XS U 140 t, XS U 124 DC-6-2-3 XS U 146 PL B 072 XS U 161 DC-4-9-11 XS U 199 XS U 232 XS U 114 tt XS U 221 tt XS U 172 ff XS U 200 XS U 123 XS U 234 t, XS U 115 tt XS U 274 XS U 164 433 DIABLO MEMBER (DIABLO CANYON)

Location Indicator Azimuth Location Indicator Azimuth

DC-6-18-20 XS U 197 DC-7b-2-6 XS U 118 ft XS U 225 tf XS U 145 fl XS U 212 DC-7b-14-17 XS U 315 11 XS U 206 Tt XS U 292 ff XS U 231 ff XS U 302

ft XS U 238 11 XS U 314

ft XS U 189 tt XS U 238

II XS U 198 11 XS U 241

ft XS U 216 If XS U 300

II XS U 215 tt XS U 158

II XS U 169 It XS U 248 It XS U 151 DC-7b-26 XS U 121 tt XS U 184 DC-7b-29-32 XS U 197

ft XS U 174 tt XS U 165

It XS U 162 tf XS U 181

tt XS U 191 tt XS U 122

XS U 235 ft XS U 169

XS U 197 It XS U 128

tl XS U 234 tf XS U 011

ff XS U 182 tf XS U 145 DC-6-20-21 XS U 158 DC-7b-35-37 XS U 273

ft XS U 192 't XS U 261

ft XS U 160 It XS U 298

ft XS U 174 tt XS U 255

DC-6-29-30 XS U 138 ft XS U 261

ff XS U 094 11 XS U 255

XS U 118 ft XS U 269 XS U 084 DC-7b-48-50 XS U 249 XS U 029 tt XS U 281

XS U 055 11 XS U 235

XS U 048 It XS U 290

ft XS U 056 'p XS U 264

DC-7a-15 XS U 215 ft XS U 237

DC-7a-18 PL B 240 t' XS U 185

DC-7a-38-40 XS U 286 ft XS U 228 ” XS U 265 11 XS U 212 ” XS U 306 DC-7b-95-97 XS U 232

ft XS U 300 tt XS U 281

DC-7b-2-6 XS U 119 tt XS U 289 ” XS U 171 DC-7b-84 XS U 309 152 DC-7b-98-100 XS U 102 085 ft XS U 111 126 If XS U 135 157 DC-7b-120-123 XS U 254 211 f I XS U 216 122 ft XS U 256 434 DIABLO MEMBER (DIABLO CANYON)

Location Indicator Azimuth Location Indicator Azimuth

DC-7b-120-123 XS U 298 DC-7b-155-158 XS U 043

DC-7b-133-135 XS U 112 ff XS U 068

XS U 166 It XS U 064

XS U 055 tt XS U rt 271 XS U 102 DC-7b-176 XS U 067 DC-7b-146-148 XS U 109 DC-7b-186-187 XS U 099 ft XS U 102 XS U 124 ft XS U 121 tt XS U 015 fl XS U 138 DC-7b-200-201 XS U 183 It XS U 118 XS U 166

DC-7b-153-155 XS U 069 tt XS U 193 ” XS U 038 DC-7b-207 XS U 262 ” XS U 089 DC-7b-212-214 XS U 133 ft XS U 076 If XS U 118 ft XS U 054 It XS U 049

DC-7b-155-158 XS U 251 tt ” XS U 113 XS U 098 XS U 092 ” XS U 055 XS U 032

OLLA MEMBER (BLACKWOOD BASIN WASH, "C" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

BW-m-6 XS U 213 BW-m-8 U 199 XS ft XS U 208 tt U 238 XS 11 XS U 224 tt XS U 231 BW-m-7 (PC1) XS U 200 ft XS U 222 XS U 229 XS U 238 XS U 228 BW-m-9 (PC1) ,XS U 231 ” XS U 201 ” XS U 189 XS U 160 XS U 219 ” XS U 235 tt XS U 189 BW-m-7 (PC2) XS U 220 BW-m-9 (PC2) XS U 204 XS U 211 XS U 198 tt XS U 219 ” XS U 242 ft XS U 226 XS U 219 1/ XS U 212 XS U 231 BW-m-8 XS U 128 BW-m-10 XS U 137 ft ” XS U 103 XS U 116 XS U 142 ” XS U 120 XS U 220 XS U 071 435 OLLA MEMBER (LAYER CAKE WASH, "C" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

LC-1-12-14 XS U 204 LC-1-182-186 XS U tt st 163 XS U 253 XS U 206 XS U 180 XS U 193 LC-1-25-28 XS U 171 t, XS U 162 XS U 168 XS U 167 tt XS U 138 tt XS U 158 XS U 168 XS U 196 If XS U 138 XS U 217 LC-1-98 XS U 161 tt XS U 224 XS U 126 XS U 217 LC-1-102-103 XS U 166 LC-1-192 U t, XS 160 XS U 151 LC-1-200 PL B 177 LC-1-103-104 XS U 194 PL B 172 XS U 153 LC-1-238-241 tt GC B 142 XS U 135 ft PL B 135 LC-1-104-106 XS U 227 PL B tt 149 XS U 173 RL U 145 t, XS U 222 LC-1-245-247 XS U 188 LC-1-121 XS U 126 XS U 162 LC-1-136 XS U 287 XS U 200 LC-1-141-145 XS U 160 XS U 220 It XS U 091 XS U 183 tt XS U 211 ft XS U 167 t, XS U 145 LC-1-248-250 XS U 181 t, XS U 193 XS U 196 It XS U 178 XS U 194 91 XS U 191 LC-1-261-263 XS U 183 LC-1-176-178 PL B 183 XS U 154 PL B 197 XS U 154 /I XS U 092 LC-1-265-266 XS U 213 tt XS U 072 XS U 178 tt XS 'U 075 XS U 182 LC-1-182-186 XS U 187 XS U 154

tt . XS U 152 436 OLLA MEMBER (OLLA WASH, "C" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

0W-1-3-4 XS U 146 0W-4-14-18 XS U 278 XS U 175 0W-4-13 XS U 281 XS U 247 0W-5-1-5 tt XS U 071 XS U 169 0W-5-1-5 XS U 048 0W-1-9-11 XS U 266 ff XS U 171 XS U 176 XS U 117 XS U 161 XS U 138 11 XS U 145 11 XS U 158 XS U 351 0W-5-7-9 XS U t, 215 XS U 345 XS U 175 t, XS U 231 t, XS U 047 XS U 192 XS U 045 ft XS U 248 XS U 267 t, XS U 238 st XS U 259 XS U 172 tt PL B 028 XS U 206 0W-5-19-21 XS U 309 PL B 191 XS U 008 t, PL B 193 0W-5-21-23 XS U 197 0W-2-1-2 XS U 235 XS U 127 XS U 168 XS U 076 t, XS U 199 tf XS U 062 11 PLU 183 t, XS U 088 tt PLU 178 0W-5-23-25 XS U 071 PL U 189 XS U 041 PL U 193 XS U 079 0W-3-2 XS U 162 XS U 023 0W-3-3 PL U 180 et XS U 058 t, PL U 169 0W-6-14-16 XS U 228 0W-3-5 XS U 214 XS U 240 0W-4-2-6 XS U 088 XS U 214 XS U 156 XS U 258 XS U 088 XS U 258 11 XS U 139 XS U 224 tf XS U 098 0W-6-16-18 XS U 094 tt XS U 132 XS U 098 ft XS U 113 XS U 154 t, XS U 088 0W-7-2-3 XS U 080 XS U 129 XS U 040 0W-4-14-18 XS U 267 t, XS U 065 XS U 224 0W-7-3-6 XS U 076 XS U 228 t, XS U 082 XS U 227 XS U 089 XS U 218 tt XS U 107 XS U 229 t, XS U 107 tt XS U 274 XS U 101 XS U 264 tt XS U 101 437 OLLA MEMBER (OLLA WASH, "C" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

OW-7-3-5 XS U 085 tt XS U 062 It XS U 126 OW-8-5-7 XS U 212 OW-8-2-4 XS U 224 It XS U 077 If XS U 196

OLLA MEMBER (SANDSTONE CANYON & VICINITY, "C" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

XS U 096 SS-m-9 XS U 205 XS U 081 XS U 205 XS U 126 f XS U 245 XS U 098 It XS U 182 XS U 103 XS U 165 XS U 089 SS-m-10 XS U 084 SS-m-7 XS U 173 XS U 116 It XS U 137 PL B 154 ft XS U 173 ft PL B 139 XS U 133

OLLA MEMBER (ARROYO TAPIADO, "C" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

AT-m-2 XS U 136 ” XS U 154 XS U 149 ” XS U 168 XS U 129 XS U 147 If XS U 209 AT-m-4 XS U 262 tt XS U 199 ” XS U 197 't XS U 172 ” XS U 202 tt XS U 179 ” XS U 231 XS U 171 AT-m-5 XS U 037 It XS U 202 ” XS U 023 ft XS U 181 ” XS U 040 AT-m-3 XS U 117 XS U 029 tt XS U 124 AT-m-6 XS U 148 XS U 176 ” XS U 185 ft XS U 189 ” XS U 166 ft XS U 163 XS U 166 XS U 153 AT-m-7 XS U 191 XS U 169 XS U 209 ft XS U 174 XS U 197 XS U 157 XS U 159 438 OLLA MEMBER (BLACKWOOD BASIN WASH, "L" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

BW-m-5 XS U 115 BW-m-5 XS U 098

OLLA MEMBER (SANDSTONE CANYON & VICINITY, "L" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

FC-m-8 XS U 053 58-m-1 XS U 077 tt XS U 039 tt XS U 078 tt XS U 187 tt XS U 032 XS U 072 SS-m-2 XS U 079 XS U 042 tt XS U 101 XS U 168 XS U 125 XS U 148 XS U 096 11 XS U 084 tt XS U 111 XS U 106 XS U 114 XS U 080 tt XS U 093 11 XS U 146 XS U 098 XS U 145 XS U 044 FC-m-9 XS U 097 XS U 055 11 XS U 189 XS U 124 11 XS U 111 XS U 096 XS U 083 tt XS U 115 XS U 102 XS U 123 FC-m-11 XS U 112 SS-m-3 XS U 149 FC-m-12 XS U 087 tt XS U 097 XS U 125 tt XS U 072 XS U 121 tt XS U 077 ft XS U 174 tt XS U 071 XS U 135 SS-m-4 XS U 051 XS U 155 tt XS U 220 tt XS U 092 tt XS U 058 FC-m-13 XS U 132 XS U 086 XS U 137 tt XS U 158 11 XS U 076 XS U 062 XS U 118 11 XS U 057 XS U 158 11 XS U 045 XS U 129 11 XS U 124 XS U 007 tt XS U 048 XS U 136 SS-m-5 XS U 191 FC-m-14 XS U 139 XS U 137 XS U 136 XS U 031 XS U 139 XS U 121 XS U 159 XS U 172 SS-m-1 XS U 353 SS-m-6 XS U 129 XS U 058 SS-m-8 XS U 070 439 OLLA MEMBER (LAYER CAKE WASH, "L" SUITE)

Location Indicator Azimuth Location Indicator Azimuth

LC-1-46-50 XS U 008 LC-1-119 XS U 211 XS U 338 LC-1-125 XS U 062 XS U 010 tt XS U 109 XS U 143 XS U 119 LC-1-81 XS U 154 LC-1-126 XS U 163 LC-1-84-85 XS U 096 LC-1-128-130 XS U 115 XS U 103 XS U 128 XS U 102 LC-1-165-168 XS U 091 LC-1-88-91 XS U 128 XS U 089 XS U 064 tt XS U 036 XS U 054 ft XS U 117 XS U 061 ft XS U 118 LC-1-90-95 XS U 041 LC-1-170-171 XS U 048 If XS U 081 XS U 109 XS U 053 LC-1-172 XS U 109 XS U 068 LC-1-207-209 XS U 055 XS U 062 XS U 085 LC-1-95-97 XS U 072 XS U 058 XS U 106 LC-1-211-213 XS U 037 XS U 143 XS U 074 LC-1-110-113 XS U 087 tf XS U 034 XS U 069 XS U 041 XS U 056 LC-1-251 XS U 180 LC-1-114 XS U 112 APPENDIX III

SANDSTONE PETROGRAPHIC DATA

Petrographic data from point counts of 29 thin sections of sandstones are presented below as raw counts and as computed percentages. Categories are mostly standard and self-explanatory; exceptions are discussed below.

1. Subdivision of volcanic lithic grains is based on unpublished notes of W. R. Dickinson (1985, written communication).

2. Carbonate sedimentary lithic grains are counted as accessory rather than framework grains, because they are strongly subject to dissolution.

3. Significant quantities of opaque "grunge" were observed in many thin sections. Under reflected light, these patches are typically yellowish to reddish, and they probably consist of a mixture of clay matrix and hematite or limonite cement.

4. Sandstones are classified on the basis of similarities to two end-member compositions, as discussed in the text: C, "C" suite; L,

"L" suite; M(C), mixed with stronger affinities for the "C" suite;

M(L), mixed with stronger affinities for the "L" suite.

440 441 RAW COUNTS

CC-m-13 CT-la-14.5 CT-2-26 CT-2-78

G (framework grains) 426 407 448 431 Qm (monocryst. quartz) 187 206 195 274 Op (polycryst. quartz) 0 4 15 12 Q (Qin + Qp + chert) 187 210 210 299 P (plagioclase) 193 165 182 35 K (K-feldspar) 41 32 47 51 F (P + K) 234 197 229 86 Lv (volcanic lithics) 1 0 0 10 microlitic/lathwork 0 0 0 5 equant felsic 0 0 0 1 vitric 0 0 0 0 hypabyssal igneous 1 0 0 4 Ls (sedimentary lithics) 0 0 0 21 argillaceous 0 0 0 8 chert 0 0 0 13 Lm (metamorphic lithics) 0 0 0 0 Li (indeterminate) 3 0 9 28 L (total lithics - chert) 5 0 9 46 Lt (total lithics + Qp) 5 4 24 71

A (accessory grains) 85 96 111 20 fossil 0 0 0 0 Ls carbonate 0 0 0 4 Mb (biotite) 79 84 103 5 Mm (muscovite) 5 12 7 0 H (heavy minerals) 1 0 1 11 matrix 24 16 0 29 Opaque "grunge" 1 28 3 7

Cc (carbonate cement) 160 185 132 189 phi (porosity) 17 3 1 32 Cc + phi 177 188 133 221

T (total counts) 712 735 695 704 recycled quartz overgrowths reworked foraminifers 442

DC-7b-209 EK-1-3.5 EK-1-18 EK-1-166-'L'

G (framework grains) 467 442 409 429 Gm (monocryst. quartz) 296 208 233 195 Op (polycryst. quartz) 13 9 12 5 Q (Qm 4- Op + chert) 324 217 256 200 P (plagioclase) 42 170 38 178 K (K-feldspar) 40 53 49 45 F (P + K) 82 223 87 223 Lv (volcanic lithics) 2 0 23 3 microlitic/lathwork 0 0 14 1 equant felsic 2 0 6 0 vitric 0 0 2 0 hypabyssal igneous 0 0 1 2 Ls (sedimentary lithics) 23 0 12 0 argillaceous 8 0 1 0 chert 15 0 11 0 Lm (metamorphic lithics) 1 0 0 0 Li (indeterminate) 50 2 42 3 L (total lithics - chert) 61 2 66 6 Lt (total lithics + Qp) 89 11 89 11

A (accessory grains) 26 77 15 200 fossil 1 0 0 0 Ls carbonate 18 0 8 0 Mb (biotite) 2 68 2 188 Mm (muscovite) 1 7 2 12 H (heavy minerals) 4 2 3 0

matrix 0 68 21 74 opaque "grunge" 2 32 1 16

Cc (carbonate cement) 258 21 105 2 phi (porosity) 11 34 5 133 Cc + phi 269 55 110 135

T (total counts) 764 674 556 854

recycled quartz overgrowths X reworked foraminifers X 443 EK-2-3 FC-1-26 FC-2-110 FC-3-17

G (framework grains) 443 435 450 446 Qm (monocryst. quartz) 213 226 209 232 Qp (polycryst. quartz) 13 12 11 12 Q (QM + Op + chert) 242 247 227 264 P (plagioclase) 71 56 120 60 K (K-feldspar) 63 40 58 37 F (P + K) 134 96 178 97 Lv (volcanic lithics) 1 32 6 12 microlitic/lathwork 0 14 1 1 equant felsic 1 15 4 7 vitric 0 1 0 0 hypabyssal igneous 0 2 1 4 Ls (sedimentary lithics) 21 18 13 32 argillaceous 5 9 6 12 chert 16 9 7 20 Lm (metamorphic lithics) 0 0 6 0 Li (indeterminate) 61 51 27 61 L (total lithics - chert) 67 92 45 85 Lt (total lithics + Qp) 96 113 63 117

A (accessory grains) 3 12 14 10 fossil 0 0 0 0 Ls carbonate 0 8 9 6 Mb (biotite) 0 3 4 4 Mm (muscovite) 2 1 0 0 H (heavy minerals) 1 0 1 0

matrix 0 8 2 5 opaque "grunge" 10 10 0 33

Cc (carbonate cement) 0 100 140 3 phi (porosity) 109 19 4 194 Cc + phi 109 119 144 197

T (total counts) 565 584 610 691

recycled quartz overgrowths reworked foraminifers 444 FC-4a-10 FC-4b-48 FC-5-50 FC-7-110 G (framework grains) 431 409 455 Qm (monocryst. quartz) 407 229 222 238 205 Qp (polycryst. quartz) 13 6 16 8 (Qui Qp + chert) 0 + 253 242 272 229 P (plagioclase) 80 60 49 53 K (K-feldspar) 46 56 48 47 F (P + K) 126 116 97 100 Lv (volcanic lithics) 13 18 54 26 microlitic/lathwork 6 8 22 12 equant felsic 6 9 22 9 vitric 0 0 3 0 hypabyssal igneous 1 1 1 5 Ls (sedimentary lithics) 16 17 20 26 argillaceous 4 3 2 10 chert 11 14 18 16 Lm (metamorphic lithics) 0 0 6 1 Li (indeterminate) 35 30 24 41 L (total lithics - chart) 52 51 86 78 Lt (total lithics + Qp) 76 71 120 102

A (accessory grains) 13 23 9 8 fossil 0 3 0 0 Ls carbonate 2 5 4 4 Mb (biotite) 7 3 1 4 Mm (muscovite) 0 1 1 0 H (heavy minerals) 4 11 3 0 matrix 18 2 64 9 opaque "grunge" 0 0 7 6

Cc (carbonate cement) 12 185 31 13 phi (porosity) 142 14 21 86 Cc + phi 154 199 52 99

T (total counts) 616 633 587 529 recycled quartz overgrowths reworked foraminifers 445 FC-m-8-'L' LC-1-14 LC-1-26 NF-1-19

G (framework grains) 427 428 437 409 0m (monocryst. quartz) 244 246 243 217 Qp (polycryst. quartz) 4 8 11 12 0 (Qm + Qp + chert) 248 262 268 234 P (plagioclase) 132 87 21 87 K (K-feldspar) 26 40 67 40 F (P + K) 158 127 88 127 Lv (volcanic lithics) 4 7 15 7 microlitic/lathwork 0 1 11 3 equant felsic 0 2 4 4 vitric 0 0 0 0 hypabyssal igneous 4 4 0 0 Ls (sedimentary lithics) 0 14 20 7 argillaceous 0 6 6 2 chert 0 8 14 5 Lm (metamorphic lithics) 6 1 2 0 Li (indeterminate) 11 25 58 39 L (total lithics - chert) 21 39 81 48 Lt (total lithics + Qp) 25 55 106 65

A (accessory grains) 90 19 26 22 fossil 0 0 0 0 Ls carbonate 0 0 17 1 Mb (biotite) 73 15 7 17 Mm (muscovite) 15 2 1 2 H (heavy minerals) 2 2 1 2 matrix 21 30 53 20 opaque "grunge" 9 4 14 10

Cc (carbonate cement) 1 12 20 85 phi (porosity) 137 97 102 42 Cc + phi 138 109 122 127

T (total counts) 685 590 652 588 recycled quartz overgrowths recycled foraminifers 446

NF-1-92(N) NF-1-194 NF-4-47 SM-1-140

G (framework grains) 420 453 428 413 Om (monocryst. quartz) 204 294 223 214 Qp (polycryst. quartz) 2 17 16 10 Q (Gin + Qp + chert) 206 318 246 224 P (plagioclase) 181 53 72 134 K (K-feldspar) 31 47 51 51 F (P + K) 212 100 123 185 Lv (volcanic lithics) 1 10 12 0 microlitic/lathwork 1 2 6 0 equant felsic 0 4 2 0 vitric 0 0 0 0 hypabyssal igneous 0 4 4 0 Ls (sedimentary lithics) 0 8 7 1 argillaceous 0 1 0 1 chert 0 7 7 0 Lm (metamorphic lithics) 0 0 0 0 3 Li (indeterminate) 1 24 47 L (total lithics - chert) 2 35 59 4 14 Lt (total lithics 4 Op) 4 59 89 68 A (accessory grains) 64 11 2 0 0 0 0 fossil 0 Ls carbonate 0 5 0 Mb (biotite) 59 2 2 58 7 Mm (muscovite) 5 2 0 3 H (heavy minerals) 0 2 0 0 matrix 0 16 19 opaque "grunge" 18 4 10 8 246 Cc (carbonate cement) 117 211 0 1 phi (porosity) 4 1 81 247 Cc + phi 121 212 81 736 T (total counts) 623 696 540

recycled quartz overgrowths X reworked foraminifers 447 SM-2-21 SM-2-54 SM-2a-7.5 SM-3-57 (framework G grains) 481 406 398 403 Gm (monocryst. quartz) 154 160 162 177 Qp (polycryst. quartz) 36 8 4 4 Q (Qm + Qp + chert) 190 168 166 181 P (plagioclase) 206 174 179 164 K (K-feldspar) 84 58 52 56 F (P + K) 290 232 231 220 Lv (volcanic lithics) 0 1 0 2 microlitic/lathwork 0 0 0 0 equant felsic 0 1 0 2 vitric 0 0 0 0 hypabyssal igneous 0 0 0 0 Ls (sedimentary lithics) 0 0 0 0 argillaceous 0 0 0 0 chert 0 0 0 0 Lm (metamorphic lithics) 1 0 0 0 Li (indeterminate) 0 5 1 0 L (total lithics - chert) 1 6 1 2 Lt (total lithics + Qp) 37 14 5 6

A (accessory grains) 150 54 55 62 fossil 0 0 0 0 Ls carbonate 0 0 0 0 Mb (biotite) 143 53 54 56 Mm (muscovite) 3 1 1 5 H (heavy minerals) 4 0 0 1 matrix 96 4 11 11 opaque "grunge" 7 15 0 10

Cc (carbonate cement) 229 107 9 88 phi (porosity) 10 19 50 34 Cc + phi 239 126 59 122

T (total counts) 972 605 523 608 recycled quartz overgrowths reworked foraminifers 448 SM -m -5

G (framework grains) 442 Om (monocryst. quartz) 156 Op (polycryst. quartz) 3 Q (Qm + Qp + chert) 160 P (plagioclase) 173 K (K-feldspar) 102 F (P + K) 275 Lv (volcanic lithics) 1 microlitic/lathwork 0 equant felsic 0 vitric 0 hypabyssal igneous 1 Ls (sedimentary lithics) 1 argillaceous 0 chert 1 Lm (metamorphic lithics) 1 Li (indeterminate) 5 L (total lithics - chert) 7 Lt (total lithics + Qp) 11

A (accessory grains) 85 fossil 0 Ls carbonate 0 Mb (biotite) 80 Mm (muscovite) 3 H (heavy minerals) 2

matrix 63 opaque "grunge" 16

Cc (carbonate cement) 166 phi (porosity) 14 Cc + phi 180

T (total counts) 786

recycled quartz overgrowths reworked foraminifers 449 COMPOSITIONAL STATISTICS (PERCENTAGES)

CC-m-13 CT-la-14.5 CT-2-26 CT-2-78

Q/ G 43.9 51.6 46.9 69.4 Qin / G 43.9 50.6 43.5 63.6 F/ G 54.9 48.4 51.1 20.0 L / G 1.2 0.0 2.0 10.7 Lt / G 1.2 1.0 5.4 16.5

Om / (Qm + P + K) 44.4 51.1 46.0 76.1 P / (Qm + P + K) 45.8 40.9 42.9 9.7 K / (Qm + P + K) 9.7 7.9 11.1 14.2

chert / Q 0.0 0.0 0.0 4.3 P / F 82.5 83.8 79.5 40.7 (Mb + Mm) / (G + A) 16.4 19.1 19.7 1.1

L / (L + Q) 2.6 0.0 4.1 13.3 Lt / (Lt + Qm) 2.6 1.9 11.0 20.6 Lv / (Lv + Ls) 32.3

phi / T 2.4 0.4 0.1 4.5 (Cc + phi) / T 24.9 25.6 19.1 31.4 classification

DC-7b-209 EK-1-3.5 EK-1-18

Q / G 69.4 49.1 62.6 46.6 Qni/ G 63.4 47.1 57.0 45.5 F / G 17.6 50.5 21.3 52.0 L / G 13.1 0.5 16.1 1.4 Lt / G 19.1 2.5 21.8 2.6

Qm / (Qm + P + K) 78.3 48.3 72.8 46.7 F / (Qm + P + K) 11.1 39.4 11.9 42.6 K / (Qm + P + K) 10.6 12.3 15.3 10.8

chert / Q 4.6 0.0 4.3 0.0 P / F 51.2 76.2 43.7 79.8 (Mb + Mm) ,/ (G + A) 0.6 14.5 0.9 31.8

L / (L + 0) 15.8 0.9 20.5 2.9 Lt / (Lt + Qm) 23.1 5.0 27.6 5.3 Lv / (Lv + Ls) 8.0 65.7

phi / T 1.4 5.0 0.9 15.6 (Cc + phi) / T 35.2 8.2 19.8 15.8

classification 450 EK-2-3 FC-1-26 FC-2-110 FC-3-17

Q / G 54.6 56.8 50.4 59.2 Qrn / G 48.1 52.0 46.4 52.0 FIG 30.2 22.1 39.6 21.7 L/ G 15.1 21.1 10.0 19.1 Lt / G 21.7 26.0 14.0 26.2

Qm / (0m + P + K) 61.4 70.2 54.0 70.5 P / (Qm + P + K) 20.5 17.4 31.0 18.2 K / (Qm + P + K) 18.2 12.4 15.0 11.2 chert / Q 6.6 3.6 3.1 7.6 P / F 53.0 58.3 67.4 61.9 (Mb + Mm) / (G + A) 0.4 1.1 0.9 0.9

L / (L + Q) 21.7 27.1 16.5 24.3 Lt / (Lt + Qm) 31.1 33.3 23.2 33.5 Lv / (Lv + Ls) 4.5 64.0 31.6 27.3 phi / T 19.3 3.3 0.7 28.1 (Cc + phi) / T 19.3 20.4 23.6 28.5 classification C C M(C)

FC-4a-10 FC-4b-48 FC-5-50 FC-7-110

Q/ G 58.7 59.2 59.8 56.3 Qin/ G 53.1 54.3 52.3 50.4 F / G 29.2 28.4 21.3 24.6 L / G 12.1 12.5 18.9 19.2 Lt / G 17.6 17.4 26.4 25.1 gm / (Qm + p + K) 64.5 65.7 71.0 67.2 P / (Qm + p + K) 22.5 17.8 14.6 17.4 K / (Qm + p K) 13.0 16.6 14.3 15.4 chert / Q 4.3 5.8 6.6 7.0 P / F 63.5 51.7 50.5 53.0 (Mb + Mm) / (G + A) 1.6 0.9 0.4 1.0

L (L + Q) 17.0 17.4 24.0 25.4 Lt / (Lt + Qm) 24.9 24.2 33.5 33.2 Lv / (Lv + Ls) 44.8 51.4 73.0 50.0 phi / T 23.1 2.2 3.6 16.3 (Cc + phi) / T 25.0 31.4 8.9 18.7 classification 451 FC-m-8-'1,' LC-1-14 LC-1-26 NF-1-19

Q / G 58.1 61.2 61.3 57.2 Gm / G 57.1 57.5 55.6 53.1 F / G 37.0 29.7 20.1 31.1 L / G 4.9 9.1 18.5 11.7 Lt / G 5.9 12.9 24.3 15.9

Qui / (Om + P + K) 60.7 66.0 73.4 63.1 P / (Qui P + K) 32.8 23.3 6.3 25.3 K / (0m + P + K) 6.5 10.7 20.2 11.6 chert / Q 0.0 3.1 5.2 2.1 P / F 83.5 68.5 23.9 68.5 (Mb + Mm) / (G + A) 17.0 3.8 1.7 4.4

L / (L + Q) 7.8 13.0 23.2 17.0 Lt / (Lt + Qui) 9.3 18.3 30.4 23.0 Lv / (Lv + Ls) 33.3 42.9 50.0 phi / T 20.0 16.4 15.6 7.1 (Cc + phi) / T 20.1 18.5 18.7 21.6 classification M(L) M(C)

NF-1-92(n) NF-1-194 NF-4-47 SM-1-140

Q/ G 49.0 70.2 57.5 54.2 Qm / G 48.6 64.9 52.1 51.8 F/ G 50.5 22.1 28,7 44.8 L/ G 0.5 7.7 13.8 1.0 Lt / G 1.0 13.0 19.2 3.4

Om / (Qui + P + K) 49.0 74.6 64.5 53.6 P / (Om + P + K) 43.5 13.5 20.8 33.6 X / (Om + P + K) 7.5 11.9 14.7 12.8

chert / Q 0.0 2.2 2.8 0.0 P / F 85.4 53.0 58.5 72.4 (Mb + Mm) / (G + A) 13.2 0.9 0.5 13.5

L / (L + 0) 1.0 9.9 19.3 1.8 Lt / (Lt + Qui) 1.9 16.7 26.9 6.1 Lv / (Lv + Ls) 55.5 63.2 phi / T 0.6 0.1 15.0 0.1 (Cc + phi) / T 19.4 30.5 15.0 33.6

Classification 452 SM-2-21 SM-2-54 SM-2a-7.5 SM-3-57

/ G 39.5 41.4 41.7 44.9 Gm / 32.0 39.4 40.7 43.9 F/ G 60.3 57.1 58.0 54.6 L / G 0.2 1.5 0.3 0.5 Lt / G 7.7 3.4 1.3 1.5

Gm / + P + K) 34.7 40.8 41.2 44.6 P / (Qm + P + K) 46.4 44.4 45.5 41.3 K / (Qm P K) 18.9 14.8 13.2 14.1 chert / Q 0.0 0.0 0.0 0.0 P / F 71.0 75.0 77.5 74.5 (Mb + Mm) / (G + A) 23.1 11.7 12.1 13.1

L / (L + Q) 0.5 3.4 0.6 5.3 Lt / (Lt + Qm) 19.4 8.0 3.0 3.3 Lv / (Lv + Ls)

phi / T 1.0 3.1 9.6 5.6 (Cc + phi) / T 24.6 20.8 11.3 20.1

classification

SM -m -5

Q / G 36.2 Gm / G 35.3 F/ G 62.2 L / G 1.6 Lt / G 2.5

Gm / (Qm + P + K) 36.2 P / (Qui + P + K) 40.1 K / (Qui + P + K) 23.7

chert / Q 0.6 P / F 62.9 (Mb + Mm) / (G + A) 15.7

L / (L + Q) 4.2 Lt / (Lt + Om) 6.6 Lv / (Lv + Ls)

phi / T 1.8 (Cc + phi) / T 22.9

classification APPENDIX IV. DATA FOR PLATES 1-4

Tie points and piercing points for fault reconstructions, Plate 1

1. Early Miocene Pinnacles (1A) and Neenach (1B) volcanics

offset -315 km across northern segment of San Andreas fault (Matthews 1976).

2. Middle Miocene Caliente Formation (2A) and Mint Canyon

Formation (2B) offset -60 km across San Gabriel Fault; latter tied to

probable source area in Chocolate Mountains (2C) offset -240 km by

southern segment of San Andreas fault (Ehlig et al. 1975; Crowell 1975a).

3. Matching Triassic plutons at Liebre Mountains (3A) and Mill

Creek (3B) offset 160 + 20 km across southern segment of San Andreas

fault (Matti et al. 1985, 1986)

4. Late Pliocene or Pliocene San Timoteo Formation offset minimum of 7 km across Banning fault (4A and 4B) (Allen 1957). Middle

Miocene Coachella Conglomerate at Whitewater Canyon (4B) offset -215 km from possible source area in Cargo Muchacho Mountains (4D) (Peterson

1975) or 180+20 km from possible source area in southern Chocolate

Mountains (46) (Dillon 1975).

5. Similar metasedimentary rocks in El Paso Mountains (5A) and

Pilot Knob Valley (5B) offset -60 km across Garlock fault; other features suggest similar offset (Smith and Ketner 1970).

453 454

6. Matching plutonic contacts and intersections of faults with

Santa Rosa cataclasite zone offset -24 km (Sharp 1967) to 28 km (Bartholomew 1970) across San Jacinto fault zone.

7. Projected Paleocene shoreline trends on SE margin of Los

Angeles basin offset -40 km across Elsinore fault (Sage 1973; Crowell and Ramirez 1979).

8. Basement rock types offset 10-20 km in fault slivers of Agua

Blanca fault (Allen et al. 1960).

9. Magdalena Fan (8A) with sandy turbidites deposited from 14.5 to 13 Ma (+0.5 Ma on biostratigraphic ages); offset 500-700 km from probable source area at incipient mouth of Gulf of California (Yeats and Haq 1981). 455 Vectors of local relative finite motions between North American and Pacific plates

4.5-0 Ma 10-4.5 Ma 21-10 Ma Average rates (Atwater and 55 mm/a 40 mm/a 13 mm/a Molnar 1973)

Azimuths (Blake et al. 1978)

36 N, 121.5 W 321.2 328.0 339.0

33 N, 119 W 318.6 325.7 335.0

26 N, 112 W 311.7 319.5 323.9 456 Additional sources for Plate 1 Faults, onshore California U.S.G.S. and Calif. Div. Mines Geol., 1966 Weldon and Humphreys 1986 Crouch, Bachman, and Shay, 1984 Bird and Rosenstock 1984

Faults, offshore northern California Crouch, Bachman, and Shay 1984 Bachman, Underwood, and Menach 1984 Blake et al. 1978 Silver and Normark 1978

Faults, offshore southern California Weldon and Humphreys 1986 Bird and Rosenstock 1984 Howell et al. 1974

Westward translation of inboard California north of Garlock fault Garfunkel 1974

Clockwise rotated domains, Transverse Ranges Luyendyk et al. 1980

Counterclockwise rotation of outboard southern California Weldon and Humphreys 1986

Salton Trough Plate 2

Faults, Pacific, offshore northern Baja California Krause 1965

Faults, Pacific, offshore southern Baja California Spencer and Normark 1979

Magdalena Fan Yeats and Haq 1981

Faults and bathymetry, Gulf of California Henyey and Bischoff 1973 Bischoff and Henyey 1974 Bischoff and Niemetz 1980 Niemetz and Bischoff 1981 457 anomalies, Magnetic mouth of Gulf and offshore southern Baja Ness et al. 1981 Curray et al. 1982 (see Chase et al. 1970, Klitgord and Mammerickx 1982 for alternative interpretations)

Colima, Zocoalco, and Chapala Grabens Allan 1985, 1986 458 Sources for Plate 2

Faults Jennings 1967 Rogers 1965, 1967 Strand 1962 Gastil, Phillips, and Allison 1971 Sharp and Clark 1972 Colletta and Ortlieb, 1984

Antithetic faults Woodard 1963 (Fish Creek-Vallecito) Fuis and Kohler 1984 (southeast of Salton Sea)

Geodetic deep slip rates Savage et al. 1979

Fold axes Dibblee 1984 (Borrego Badlands, San Felipe Hills, Superstition Hills) Sharp and Clark 1972 (Borrego Badlands) Babcock 1974 (Durmid Hill) Proctor 1968 (Indio Hills) Sylvester and Smith 1976 (Mecca Hills)

Spreading centers Fuis and Kohler 1984

Volcanos Fuis and Kohler 1984

Basement escarpments Fuis and Kohler 1984

Rotational domains Johnson et al. 1983 Mace 1981 Bogen and Seeber 1986

Shoreline of Lake Cahuilla Wilke 1978 459 Abbreviated localities in Plate 3

AC - Agua Caliente County MW - Mortero Wash Park NF - North Fork, Fish Creek AH - Arroyo Hueso Wash AS - Arroyo Salada NMW - North Mortero Wash ASD - Arroyo Seco del Diablo OW - 011a Wash AT - Arroyo Tapiado PD - Pinyon Dropoff BC - Buttes Canyon PG - Painted Gorge BMW - Borrego Mountain Wash PMR - Pinyon Mountain Road BSW - Borrego Sink Wash PST - Palm Spring Turnoff BWC - Bow Willow Campground PVW - Palo Verde Wash CAT - Cut-Across Trail PW - Pinto Wash CBO - Carrizo Badlands RW - Rainbow Wash Overlook SC - Sandstone Canyon CC - Calcite Canyon Scenic SFW - San Felipe Wash Area SMG - Split Mountain Gorge CSN - Canyon Sin Nombre SW - Short Wash DCC - Dos Cabezas Primitive TGC - Tamarisk Grove Campground Campground DD - Diablo Dropoff TT - Thimble Trail DV - Davies Valley TW - Tule Wash EW - Ella Wash TAW - Tarantula Wash FCC - Fish Creek Primitive USG - U. S. Gypsum quarry Campground VC - Vallecito Creek FCW - Fish Creek Wash VM - Vista del Malpais FP - Fonts Point VSP - Vallecito Stage Station FPW - Fonts Point Wash County Park HM - Hills of the Moon Wash YPC - Yaqui Pass Primitive IHM - Imperial Highway Campground Monument YW - Yuha Wash JW - June Wash 5P - Five Palms JJW - Jojoba Wash 17P - Seventeen Palms MPS - Mountain Palm Springs

Sources for Plate 3

Base map Weber 1959 Morton 1966 Lindsay and Lindsay 1978

Geology Dibblee 1984 Weber 1959, after Dibblee 1954 Morton 1966, after Dibblee 1954 Hoover 1965 field reconnaissance by author Plate 4 460 Sources for Plate 4

Fish Creek and Vallecito basins Woodard 1963

Carrizo Impact Area Kerr 1982

Vallecito Basin P. Remeika 1985, written communication

Split Mountain, SE flank Vallecitos, Sweeney Pass quadrangle field mapping by author

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