. . THE GEOLOGICAL SOCIETY • OF AMERICA®
Field Guide 11
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edited by Ernest M. Duebendorter and Eugene I. Smith
Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada
edited by Ernest M. Duebendorfer Northern Arizona University Geology Department Frier Hall Knoles Drive Flagstaff, Arizona 86011-4099 USA Eugene I. Smith Department of Geoscience University of Nevada, Las Vegas 4505 S. Maryland Parkway Las Vegas, Nevada 89154-4010 USA
Field Guide 11 3300 Penrose Place, P.O. Box 9140
Boulder, Colorado 80301-9140 USA
2008
Copyright © 2008, The Geological Society of America, Inc. (GSA). All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact the Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. Library of Congress Cataloging-in-Publication Data Field guide to plutons, volcanoes, faults, reefs, dinosaurs, and possible glaciation in selectedareas of Arizona, California, and Nevada / edited by Ernest M. Duebendorfer, Eugene I. Smith. p. cm. -- (Field guide ; 11) Includes bibliographical references. ISBN: 978-0-8137-0011-3 (pbk.) 1. Geology--Arizona. 2. Geology--California. 3. Geology--Nevada. I. Duebendorfer, Ernest M. II. Smith, Eugene I. (Eugene Irwin), 1944QE85.F54 2008 557.9--dc22 2008006898 Cover: Spectacular geology in the Lake Mead area just west of Las Vegas. The River Mountains volcanic section (foreground in Nevada) and the Wilson Ridge pluton (on the skyline to the east in Arizona) represent a linked volcanic-plutonic system separated by the Saddle Island detachment fault. The mesa is Fortification Hill capped by 5.8 m.y. old basalt. Photo by Eugene Smith, May 2006.
10 9 8 7 6 5 4 3 2 1
ii
Contents
Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. The mid-Miocene Wilson Ridge pluton and River Mountains volcanic section, Lake Mead area of Nevada and Arizona: Linking a volcanic and plutonic section . . . . . . . . . . . . 1 Denise Honn and Eugene I. Smith 2. Late Paleozoic deformation in central and southern Nevada . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 Pat Cashman, Jim Trexler, Walt Snyder, Vladimir Davydov, and Wanda Taylor 3. Active tectonics of the eastern California shear zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43 Kurt L. Frankel, Allen F. Glazner, Eric Kirby, Francis C. Monastero, Michael D. Strane, Michael E. Oskin, Jeffrey R. Unruh, J. Douglas Walker, Sridhar Anandakrishnan, John M. Bartley, Drew S. Coleman, James F. Dolan, Robert C. Finkel, Dave Greene, Andrew Kylander-Clark, Shasta Marrero, Lewis A. Owen, and Fred Phillips 4. Ediacaran and early Cambrian reefs of Esmeralda County, Nevada: Non-congruent communities within congruent ecosystems across the Neoproterozoic-Paleozoic boundary . . . . 83 Stephen M. Rowland, Lynn K. Oliver, and Melissa Hicks 5. Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 Andrew P. Barth, J. Lawford Anderson, Carl E. Jacobson, Scott R. Paterson, and Joseph L. Wooden 6. Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary, northwest Arizona: A tale of three basins, immense lacustrine-evaporite deposits, and the nascent Colorado River . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119 James E. Faulds, Keith A. Howard, and Ernest M. Duebendorfer 7. Interpretation of Pleistocene glaciation in the Spring Mountains of Nevada: Pros and Cons . . . 153 Jerry Osborn, Matthew Lachniet, and Marvin (Nick) Saines 8. Quaternary volcanism in the San Francisco Volcanic Field: Recent basaltic eruptions that profoundly impacted the northern Arizona landscape and disrupted the lives of nearby residents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173 S.L. Hanson, W. Duffield, and J. Plescia 9. The Spirit Mountain batholith and Secret Pass Canyon volcanic center: A cross-sectional view of the magmatic architecture of the uppermost crust of an extensional terrain, Colorado River, Nevada-Arizona . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 187 Nicholas P. Lang, B.J. Walker, Lily L. Claiborne, Calvin F. Miller, Richard W. Hazlett, and Matthew T. Heizler iii
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Contents
10. Devonian carbonate platform of eastern Nevada: Facies, surfaces, cycles, sequences, reefs, and catastrophic Alamo Impact Breccia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 215 John E. Warme, Jared R. Morrow, and Charles A. Sandberg 11. Dinosaurs and dunes! Sedimentology and paleontology of the Mesozoic in the Valley of Fire State Park . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249 Joshua W. Bonde, David J. Varricchio, Frankie D. Jackson, David B. Loope, and Aubrey M. Shirk
Preface
Welcome to Las Vegas! This guidebook has been prepared in conjunction with the 2008 combined Cordilleran and Rocky Mountain Sections meeting of the Geological Society of America. This volume contains background information and road logs for eleven field trips in Nevada, Arizona, and California. Southern Nevada and adjoining areas contain a rich geologic history spanning the interval from the Paleoproterozoic to the present. Las Vegas lies at or near several critical geological junctures and localities including the structural boundary between the Colorado Plateau and Basin and Range, the physiographic boundary between the Great Basin and the southern Basin and Range, the eastern margin of the Sevier foldand-thrust belt, the tectonically active Death Valley area, tilted and faulted volcanic-plutonic systems exposing the upper part of the crust, and the enigmatic “amagmatic zone.” Field trips in this volume span the geologic record from the Ediacaran (late Neoproterozoic) to the Holocene. Steve Rowland, Lynn Oliver, and Melissa Hicks will lead participants to three of the best examples of Ediacaran and Early Cambrian reefs in North America (Chapter 4). A trip led by John Warme, Jared Morrow, and Charles Sandberg (Chapter 10) examines the long-lived Devonian shallow-water carbonate platform and features a visit to the spectacular Alamo Impact Breccia. Middle Mississippian to late Permian tectonism as recorded by regional unconformities, folding, thrusting, and the stratigraphic record is the focus of a trip led by Pat Cashman, Jim Trexler, Walt Snyder, Vladimir Davydov, and Wanda Taylor (Chapter 2). Andy Barth, Lawford Anderson, Carl Jacobson, Scott Paterson, and Joe Wooden bring us into the Mesozoic with an overview of the tectonic evolution of a tilted section through the upper and middle crust of the Cretaceous Cordilleran arc (Chapter 5). Cretaceous sedimentary rocks deposited in the foredeep of the Sevier fold-and-thrust belt and their dinosaur fossils are the topic of a trip led by Joshua Bonde, David Varricchio, Frankie Jackson, David Loope, and Aubrey Shirk (Chapter 11). The Cenozoic is well represented by six different trips. Nick Lang, B.J. Walker, Lily Claiborne, Calvin F. Miller, Rick Hazlett, and Matt Heizler (Chapter 9) examine spectacular cross-section view of the Miocene Spirit Mountain batholith and a coeval, and possibly related, eruptive center (Secret Pass) in the Colorado River extensional corridor. Another volcano-plutonic complex, the River Mountains–Wilson Ridge igneous system, which was dismembered by the Saddle Island detachment fault is the destination of a trip led by Denise Honn and Gene Smith (Chapter 1). Jim Faulds, Keith Howard, and Ernie Duebendorfer examine synextensional basins that constrain the timing of the structural demarcation between the Colorado Plateau and the Basin and Range (Chapter 6). Jerry Osborn, Matthew Lachniet, and Nick Saines weigh the evidence for and against Pleistocene glaciation in the Spring Mountains of southern Nevada in Chapter 7. The cultural effects of some of the youngest volcanism in the continental United States outside the Cascades is the focus of a trip by Sarah Hanson, Wendell Duffield, and Jeffrey Plescia (Chapter 8) to the San Francisco volcanic field near Flagstaff, Arizona. Finally, Kurt Frankel and a cast of thousands bring us up to date with a look at the active tectonics of the eastern California shear zone with discussions regarding significant discrepancies between long-term slip rates and the current rate of strain accumulation along active faults (Chapter 3). With field trips ranging from old to the present, the middle crust to the surface, from tectonics to paleontology, and from volcanism to glaciation, this volume offers something for everyone. Ernest M. Duebendorfer Eugene I. Smith
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Map of the Nevada, California, Arizona, and Utah areas visited in these field trips showing locations of trips by number.
The Geological Society of America Field Guide 11 2008
The mid-Miocene Wilson Ridge pluton and River Mountains volcanic section, Lake Mead area of Nevada and Arizona: Linking a volcanic and plutonic section Denise Honn* Eugene I. Smith* Department of Geoscience, University of Nevada, Las Vegas, Nevada 89154-4010, USA
ABSTRACT This field trip will visit the River Mountains volcanic section (12.17 ± 0.02 to 13.45 ± 0.02 Ma) and Wilson Ridge pluton (13.10 ± 0.11 Ma) in southern Nevada and northwestern Arizona. Although this volcanic-plutonic system was disrupted by the Saddle Island detachment fault during Miocene crustal extension, there are convincing lithological, mineralogical, geochemical and geochronological indicators that suggest a cogenetic relationship. The trip consists of 17 stops that emphasize evidence that links the volcanic and plutonic sections. In addition we will visit the Saddle Island detachment fault at its type locality on Saddle Island. Keywords: plutonic rocks, volcanoes, Lake Mead, petrology, geochronology. The River Mountains volcanic section–Wilson Ridge pluton igneous system crops out at the northern end of the Colorado River extensional corridor, a north-south trending area of southern Nevada, western Arizona and eastern California that underwent up to 100% extension between ca. 23 and 12 Ma. In the northern part of the corridor, volcanic rocks of Tertiary age lie on Precambrian crystalline rocks and locally a thin conglomerate containing sedimentary and crystalline clasts. Paleozoic and Mesozoic sedimentary sections are missing and were probably stripped from a rising structural arch (the Kingman Arch) during late-Cretaceous, early Tertiary time (Faulds et al., 2001). The arch plunges gently to the north (~15°) and terminates against the Lake Mead fault system just north of Lake Mead. In the Colorado River extensional corridor, magmatism migrated to the north, pre-dating crustal extension by about 1 m.y. (Faulds et al., 2001).
INTRODUCTION The study of an igneous system is limited by exposure and preservation of the rock record. In most cases, only a portion of the system is exposed (i.e., volcanic or plutonic) and therefore only part of the magmatic history can be studied. Based on work done over the past 20 years, we interpret the River Mountains volcanic section of southern Nevada and the Wilson Ridge Pluton in northwestern Arizona as volcanic and plutonic segments of the same igneous system (Fig. 1). The connection between the River Mountains volcanic section and the Wilson Ridge pluton is based on structure, lithology, mineralogy, geochemistry, and geochronology. This field trip will visit both the River Mountains and Wilson Ridge and will emphasize links between the volcanic and plutonic sections. *
[email protected],
[email protected] Honn, D., and Smith, E.I., 2008, The mid-Miocene Wilson Ridge pluton and River Mountains volcanic section, Lake Mead area of Nevada and Arizona: Linking a volcanic and plutonic section, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 1–20, doi: 10.1130/2008.fld011(01). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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Figure 1. Geologic map of Lake Mead region. Adapted from Smith et al. (1990).
Boulder City
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Mid-Miocene Wilson Ridge pluton and River Mountains volcanic section Quaternary
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Qts - Tertiary and younger sediments Tmf - Fortification Hill basalt
Faults - dotted where concealed or inferred ball on downthrown side
Tcf - Callville Mesa volcanic rocks
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Thc - Hamblin-Cleopatra volcanic rocks
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Tgp - Unassigned volcanic rocks
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Tvu - upper Hoover Dam volcanic rocks Powerline road volcanic rocks Tpd - dacite Tpd - basalt and andesite Tbc - Boulder City Pluton Tbv - Bootleg Wash volcanic rocks Thd - Tuff of Hoover Dam and Dam Conglomerate Tbw - Boulder Wash volcanic rocks Tbr - Breccia River Mountains Stratovolcano Trs - quartz monzonite stock Tsv - andeste and dacite
Wilson Ridge Pluton Twrh - hypabyssal phase Twrm - medium grained phase Twrc - coarse grained phase Twrg - red feldspar granite phase Tid - diorite phase
Tpm - Patsy Mine volcanic rocks K-Tpp - Paint Pots intrusive rocks
Mesozoic Mz-p - Pennsylvanian through Mesozic rocks Paleozoic Precambrian
M-Pc - Precambrian through Mississippian rocks
Figure 1 (continued).
RIVER MOUNTAINS VOLCANIC SECTION The River Mountains volcanic section (12.17 ± 0.02 to 13.45 ± 0.02 Ma, 40Ar/39Ar whole-rock and mineral dates; Faulds et al., 1999) composed of mainly dacite, andesite, basalt and rhyolite is locally intruded by hypabyssal dacite plugs and a quartz monzonite stock. Smith (1982; 1984) suggested that the River Mountains are composed of at least four volcanoes that were juxtaposed by mid-Tertiary strike-slip faulting related to the leftlateral Lake Mead fault system: 1. The River Mountains stratovolcano and related satellitic dacite, rhyolite and basalt volcanoes. The stratovolcano is cored by the River Mountains quartz monzonite stock, which is surrounded by a zone of altered volcanic rocks cut by numerous dikes. The stock contains many xenoliths of basalt and dolomite. Dikes of porphyritic dacite radiate from the plug. The stock is chemically equivalent to rocks of the Wilson Ridge pluton and may represent the detached apex of one of the Wilson Ridge intrusions. Rocks above the intrusion are altered and mineralized andesite and plutonic rock cut by numerous dacite dikes that emanate from the River Mountains stock. The blue sodium amphibole, magnesio-riebeckite, occurs along
fractures and coatings on rocks of the quartz monzonite stock and surrounding altered volcanic rocks. Magnesioriebeckite is also found in fractures and thin veins in various phases of the Wilson Ridge pluton and within the Colorado River extensional corridor appears to be unique to this volcanic-plutonic system. 2. The Bootleg Wash section just north Boulder City, Nevada, composed from base to top of a section of andesite flows, volcaniclastic breccia, and flow-banded dacite flows. 3. The Red Mountain section formed by highly altered andesite and dacite flows, volcaniclastic rocks, and local granitic intrusions. On Red Mountain, andesite flows and breccia are interleaved along numerous low-angle faults. The Red Mountain section is separated from the River Mountains stratovolcano by a northwest-striking fault (probably strike slip) and may represent highly altered volcanic and plutonic rocks related to the Boulder City pluton (14.17 ± 0.6 Ma; NAVDAT (http://navdat. kgs.ku.edu/); 13.8 Ma K/Ar age reported by Anderson et al., 1972). 4. The Casino dacite just east of Railroad Pass is characterized by andesite and dacite flows and a thin moderately welded ash-flow tuff.
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WILSON RIDGE PLUTON (ABSTRACTED FROM LARSEN AND SMITH, 1990) The Wilson Ridge pluton is an epizonal to hypabyssal calcalkaline pluton that formed ca. 13.10 ± 0.11 Ma (40Ar/39Ar hornblende date; Faulds et al., 1999) during a period of mid-Miocene extension. Faulting and erosion have provided a cross section of the pluton in plan view. Geobarometric data and geologic constraints indicate the pluton has been tilted 17° to the north (Metcalf et al., 1993). The apex of the pluton, just south of Boulder Canyon (Lake Mead), Nevada, is composed of hypabyssal quartz monzonite and dacite cut by numerous dikes of rhyolite, dacite and basalt. The base of the pluton is 20 km to the south where quartz monzodiorite, monzodiorite, and diorite are in low-angle intrusive contact with Precambrian basement. The pluton was separated from comagmatic volcanic rocks in the River Mountains by movement along the Saddle Island fault system which includes the Saddle Island detachment, Hamblin Bay, and Eldorado faults (Weber and Smith, 1987; Duebendorfer et al., 1990). The age of detachment is estimated to be younger than ca. 13.5 Ma based on the inference that detachment faulting must postdate the formation of the Wilson Ridge pluton and River Mountain volcanic suite (Duebendorfer et al., 1990). The River Mountains now lie approximately 20 km to the west of the pluton. The Wilson Ridge pluton is composed of the Teakettle Pass suite consisting of foliated monzodiorite and quartz monzodiorite, unfoliated quartz monzonite, and the older Horsethief Canyon diorite. The Teakettle Pass suite comprises the main phase of the Wilson Ridge pluton (80 km2 outcrop area). The major minerals of the coarse-grained quartz monzonite (the dominant phase of the Teakettle Pass suite) are quartz (20%), orthoclase (25%), plagioclase (40%), and subhedral prismatic hornblende (60 k.y. of surface-exposure. Interestingly, it was also found that this inheritance varied with grain size (Fig. 4). Despite the high inheritance, the depth-profile approach yielded a very well constrained date of 37 ± 7 ka for the abandonment of the Q2b alluvial fan depositional surface.
Fault Offsets and Along-Strike Slip Rate Gradients Offsets of Q2b alluvial fans and related inset channels are apparent in both the southern and northern field stops (Stops 1 and 2; Figs. 1, 2, and 3), though the northern set of offsets is better defined. Here, combined constraints from a pair of inset channels and a shutter ridge yield an offset of 30 ± 5 m (Fig. 3). Channel deflections in the southern field stop (Stop 1) are complicated by multiple fault strands and fan emplacement along the fault scarp (Fig. 2). Offsets of the larger channels proved difficult to constrain, but must be 15 m in the Summit Digging location, but the total thickness and inclusion of all tuffs in the sequence at correlative outcrops varies widely. A distinctive pink to orange-pink ash-flow tuff lies within this unit and crops out prominently in the middle of the Summit Range area. This pumiceous lapilli tuff is rich in lithic clasts with compositions including propylitically altered porphyritic extrusive volcanic rocks, plutonic rocks of varying composition, quartz, and banded rhyolite (Fig. 5). It also has a large percentage by volume of pumice clasts. It reaches a thickness of several meters in outcrop in the Summit Digging area, and crops out widely on the floor of the valley, where it can be found overlying the tan and deep red lahars, debris flows, and sandstones, and in other places be found beneath them. The age of this unit is 11.8 ± 0.9 Ma (from (U-Th)/He on zircon). The next highest stratigraphic unit consists of a series of dacite domes, flows, and tuffs clustered in a 5 km2 area situated on the western and southern margins of the Summit Range. These units overlie the white tuffs that form the valley floor between the westernmost flows and the easternmost domes and overlie the tuffs on a small, elliptical, NE-trending hill 2 km south of the main diggings. Units in the domes are quite distinctive because of their coarse phenocrysts of twinned (Carlsbad habit) orthoclase that reach lengths of several centimeters. These are embedded in a fine-grained, reddish-brown matrix with biotite and small quartz grains. On the easternmost dome, there is a vertical fabric to the rock. The dome is surrounded on two sides by outcrops of lava flows with the same, only finer-grained, mineral composition as the dome. These flows exhibit abundant flow features, such as alignment of phenocrysts in flow bands, and contorted, brecciated flow fronts. Flows at the southeast corner of the dome dip 30–40° outward and show well-developed cooling column structure. The radiometric age of this dome is 11.0 ± 0.2 Ma (40Ar/39Ar). A second dome, located ~1 km southwest of the dome just described, is petrologically similar although it is highly altered,
Active tectonics of the eastern California shear zone
Summit Range
Dacite domes and associated flows and tuffs 11.7 ± 0.6 Ma to 11.02 ± 0.19 (M&W)
51
Red Rock Canyon
Tda4 air fall tuff 8.5 ± 0.13 Ma (L&B)
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Tda3 air fall tuff 8.4 ± 1.8 Ma (L&B) Tda2 air fall tuff 10.4 ± 1.6 Ma (L&B) Tdb3 basalt flow 10.5 ± 0.25 Ma (L&B)
Td4 Tdb2 basalt flow 11.3 ± 0.3 Ma (M&W)
Td3 White and green tuffs Tda1 lapilli tuff 11.8 ± 0.9 Ma (L&B) (Tuff of Dutch Cleanser Mine)
Lapilli tuff 11.8 ± 0.9 Ma (M&W)
White and green tuffs Air fall tuff 11.7 ± 0.3 Ma (M&W) Red to tan arkosic conglomerate and sandstone,olive drab to brown mudstone, limestone, and sandstone
Td2 Conglomerate, sandstone, mudrock, limestone, chert
Gray andesite porphyry 15.6 ± 0.5 Ma (M&W)
Atolia quartz monzonite 86.0 ± 1.0 Ma (M&W)
Figure 5. Schematic stratigraphic sections for the Summit Range and Red Rock Canyon areas. Important and/or dated units shown in pattern; unpatterned rocks consist of conglomerate, sandstone, mudstone, limestone, and chert. Ages are from Monastero and Walker (unpublished data), denoted M&W, and from Loomis and Burbank (1988) and Whistler and Burbank (1992), denoted L&B. Units and descriptions for the Red Rock Canyon area are from Loomis and Burbank (1988).
mostly massive, coarsely porphyritic dacite. There is pervasive argillic alteration of orthoclase and biotite phenocrysts and a general vertical fabric, which suggests flow within a vent. The heavily altered area covers ~75–100 m2 and stands 8–10 m high. Numerous examples of monolithologic breccias reminiscent of vent structures can be found on the top of the dome. Flows extend to the south and west of the dome complex and thin rapidly to 1–3 m within 1 km. The age of this dome is determined to be 11.7 ± 0.6 Ma (40Ar/39Ar). Middle to Late Miocene Stratigraphy of the Red Rock Canyon Area Approximately 30 km west-southwest of the Summit Range is the Red Rock Canyon area (Stop 4, Fig. 1). Situated at the
southwest end of the El Paso Mountains, this area is the type section for the middle to late Miocene Dove Springs Formation (Loomis, 1984). The stratigraphic succession for this interval is shown in Figure 5. Loomis (1984) divided the rocks into six units, ranging in age from 13.5 Ma to ca. 7 Ma; the Dove Springs rests disconformably on the early to middle Miocene Cudahy Camp Formation. The hiatus between the Cudahy Camp and Dove Springs Formations was determined by Loomis and Burbank (1988) to be 15.1–13.5 m.y. based on magnetostratigraphy, K-Ar, and fission track dates. The lowest part of the Dove Springs section contains conglomerate and sandstone consisting of poorly sorted, massive to crudely stratified units dominated by volcanic and plutonic clasts in an ash-rich matrix. The unit grades continuously upward into massive, dirty arkosic sandstones. The top of this unit is
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arbitrarily defined as the point where the sandstones and conglomerates become less massive and more stratified. There are no age-specific geochronology markers in this unit. Member 2 of the Dove Spring Formation (Td2) crops out extensively in the south-central part of Red Rock Canyon State Park where it is dominated by conglomerates, sandstones, and volcanic tuffs (Fig. 5). Loomis (1984) determined that the provenance of these rocks was from a source located to the south and east. The sedimentary and epiclastic rocks consist of grains of quartz, feldspars, and biotite, with varying degrees of rounding, and ash, pumice, and lithic clasts from volcanic and plutonic sources. They vary in color from deep red to tan, are intimately interleaved, and constitute between 200 and 400 m of section. In the middle of Td2, there is a pink lapilli ash-flow tuff (Tda) that is a prominent ridge-former in the Red Rock Canyon area (Fig. 5). The unit is made up of two separate flows, with a cooling break between them. The rock is rich in pumice clasts (0.5– 2.0 cm) and lithics (propylitically altered volcanics, banded rhyolites, and plutonic fragments) in an ash- and crystal-rich matrix. This unit grades northeastward to a white to light gray air-fall tuff dated at 11.8 ± 0.9 Ma (Whistler and Burbank, 1992). There are several thin (tens of centimeters) white air-fall and pumicerich tuffs below this unit that have been dated at 11.7 ± 0.6 Ma (40Ar/39Ar). Overlying the pink tuff is a continuation of the Td2type rocks described in the foregoing paragraph. Higher in the section there are two basalt flows (Tdb2 and Tdb3 of Loomis, 1984), the lower of which have been dated at 11.3 ± 0.3 Ma (40Ar/39Ar; Monastero and Walker, unpublished data). These flows are interleaved with more of the Td2 type rocks although they are grouped into higher units of the Dove Spring.
time of initiation of extension in Death Valley. This would mark the beginning of movement on the Garlock fault inasmuch as there would have been no need for such a structure prior to this time. The assumption is, therefore, that initial movement on the eastern part of the Garlock fault began ca. 15 Ma, which means that there was ~30 km of sinistral offset on the Garlock from inception of movement until 12–11 Ma. This translates to a rate of offset of between 7 and 10 mm/yr. Since that time, an additional ~35 km of offset has resulted in the present spatial relationship of the two sites, which calculates to a slip rate of 2.75–3.0 mm/yr—much slower than the rate of the earlier period. The specific cause of this dramatic decrease in offset rate on the Garlock fault is not known with certainty at present. Lonsdale (1991) and Atwater (1989) document a change in the configuration of the Pacific and North America plates ca. 12.5 Ma. Lonsdale (1991) contends that the plate offshore Baja California stopped spreading at this time, subduction ceased, and the Rivera triple junction jumped southward to the tip of the Baja peninsula. Atwater (1989) found that at this same time there was simultaneous formation of a continuous boundary between these two plates from central California to the tip of Baja. There are numerous examples of volcanic outbursts during this time in the nearby Owlshead Mountains (Calzia and Ramo, 2000), Death Valley (Thompson et al., 1993; Troxel, 1994), and Eagle Crags volcanic field (Monastero et al., 1994). It is possible, and actually highly probable, that the rate of sinistral offset on the Garlock fault varied considerably from 12 to 11 Ma until present, but such correlations are left to a more detailed understanding of the stratigraphy of the middle to upper Dove Spring Formation and the volcanic units found in the Lava Mountains (Smith et al., 2002).
Implications of Unit Correlation at Red Rock Canyon and the Summit Range
Summary
Although the relative positions of the Summit Range volcanic center and the southern Red Rock Canyon stratigraphic sequences at the time of emplacement cannot be determined with absolute certainty, it is clear that they correlate in age and lithology over a wide range of time. It is suggested that the Summit Range location was the source area for most of the volcanic components of the lower Dove Spring Formation in Red Rock Canyon. Based on thickness of the pink lapilli tuff and the fact that the Tda unit makes a transition to a pure white air-fall tuff in a north-northeasterly direction, we interpret that the two sites were more or less juxtaposed across what is now the Garlock fault when the volcanic center was active between 12 and 11 Ma. The notion of the Garlock being an intracontinental transform (Davis and Burchfiel, 1973) that acts as an accommodation zone for extension north of the fault from a relatively unextended area south of the fault, implies that movement on the Garlock could not have initiated before the onset of extension in the southwest Basin and Range. Wernicke et al. (1988) report the onset of extension in the Las Vegas Valley shear zone–Lake Mead area ca. 16 Ma, and McKenna and Hodges (1990) cite 15 Ma as the
The rocks exposed in the Summit Range and Red Rock Canyon demonstrate that at 12–11 Ma the two areas, which are now separated by ~35 km of sinistral offset on the Garlock fault, were directly opposite one another across that structure. This implies that initial motion on the Garlock fault was on the order of 10 mm/yr from ca. 15–12 Ma. If this is correct, then the later offset rate is a factor of three to four times slower, averaged over the past 11 Ma. BEDROCK EVIDENCE FOR 65 km OF DEXTRAL OFFSET ACROSS OWENS VALLEY, CALIFORNIA, SINCE 83 Ma Historic earthquakes (Beanland and Clark, 1994), paleoseismology (Bacon and Pezzopane, 2007; Lee et al., 2001b), and geodesy (Dixon et al., 2003) indicate that the modern tectonic regime of Owens Valley is dominated by right slip that accommodates a significant fraction of the relative motion between the North America and Pacific plates. However, the total magnitude and the timing of lateral slip across Owens Valley have been uncertain. Conventional wisdom has long been that net lateral
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Golden Bear and Coso Dikes Moore (1963, 1981) first recognized the Golden Bear dike and mapped it for ~15 km, from its western termination in the northern Mount Whitney quadrangle to where it disappears under Owens Valley near Independence (Fig. 6). The dike actually comprises from one to three branching granitic dikes that range from 5 to 30 m thick, and locally have thick (50 cm) cataclastic margins. The Coso dike swarm includes two groups of steeply dipping E-striking dikes that are separated by ~6 km across strike (Duffield et al., 1980; Whitmarsh, 1998). Major dikes of the Coso swarm are from 5 to 25 m thick, commonly accompanied by thinner (20%
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Figure 7. Map of quantitative properties of the Jurassic Independence dike swarm, modified from Bartley et al. (2008). Shaded areas are preCenozoic bedrock. Symbols are centered on midpoints of traverses. Dashed symbols indicate traverses where observed dikes may be Cretaceous and therefore not part of the Independence swarm (see Bartley et al., 2008, for further discussion). (A) Percent dilation by diking. Note offset of high-dilation zone between the Alabama Hills (AH) and Coso Range and the distinct northern boundary on both sides of Owens Valley where high dilations drop to zero. SH—Spangler Hills. (B) Number of dikes crossed per kilometer. The pattern resembles that in (A), but traverses in the southeastern Sierra Nevada (dashed) appear more prominent owing to much higher abundances of thin (5–30 cm) mafic dikes. Such a high proportion of thin mafic dikes is atypical of the Independence swarm and more characteristic of Cretaceous dikes in the region. (C) Kamb contour plot of poles to dikes measured on traverses. GF—Garlock fault; SAF—San Andreas fault; CA—California; NV—Nevada.
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Garlock fault is transmitted to the north (e.g., Oskin and Iriondo, 2004). Thus, it seems most likely that a significant fraction of the total 65 km of slip—perhaps as much as 50–60 km—is significantly older, possibly Late Cretaceous (Bartley et al., 2008). Laramide-age right slip along the eastern margin of the Sierra Nevada may have been linked to a south-directed extensional detachment system in the southern Sierra Nevada (Wood and Saleeby, 1997) that unroofed high-pressure rocks in latest Cretaceous–early Tertiary time. Dextral shear of this age would predate formation of the Garlock fault (Monastero et al., 1997) and thus eliminate that conflict. Locus of Slip The fault zone responsible for offset of the various markers lies between the Sierra Nevada and White-Inyo-Coso Ranges (Fig. 6). Although mostly buried by alluvium, it can be located within 100 m at Little Lake on the west side of the Coso Range, where Jurassic plutonic rocks with abundant Independence dikes are juxtaposed against a distinctive Jurassic orthogneiss that lacks such dikes (Stop 5; Figs. 1 and 8; Bartley et al., 2008). Cretaceous and Jurassic intrusive rocks on both sides of the fault in this area are cut by transpressive ductile-brittle shear zones with a persistent dextral component of shear. These shear zones, and similar Late Cretaceous dextral transpressive shear zones exposed in the Inyo (Vines, 1999) and White (Sullivan and Law, 2007) Mountains, probably belong to a family of structures that accommodated 50+ km of dextral offset in Late Cretaceous– early Tertiary time. Approaching the main fault contact at Little Lake, cataclastic microstructures increasingly overprint the ductile microstructures in the shear zones. This spatial pattern may reflect the reactivation of a Late Cretaceous shear system by the modern right-slip fault zone. Summary Several independent geologic markers indicate 65 ± 5 km of net dextral shear across Owens Valley since 83 Ma. On the order of 10 km of this displacement accumulated in the modern dextral transtension regime, and the remainder probably accumulated in Late Cretaceous–early Tertiary time. The modern tectonic regime therefore appears to reflect reactivation of a long-lived throughgoing crustal boundary. QUATERNARY TECTONISM OF THE NORTHWESTERN COSO RANGE The Coso Range is a tectonically and volcanically active region along the southeastern margin of the Sierra Nevada (“Sierran”) microplate, which moves ~13 mm/yr northwest with respect to stable North America (Argus and Gordon, 1991, 2001; Dixon et al., 1995, 2000). Northwest motion of the Sierran microplate is accommodated by distributed strike-slip and normal faulting in a 100-km-wide zone of active deformation
bordering the eastern Sierra Nevada (Fig. 1; Dokka and Travis, 1990; Unruh et al., 2003). Active crustal extension in the Coso Range primarily is driven by a releasing transfer of dextral motion from the Airport Lake fault to the Owens Valley fault, two major right-lateral strike-slip faults that form the eastern tectonic boundaries of the Sierran microplate south and north, respectively, of the Coso Range (Fig. 1; Monastero et al., 2005; Unruh et al., 2002). In detail, the Airport Lake fault zone splits into several branches at the southern end of the step-over region (Fig. 9). An eastern branch, consisting of N- to NNE-striking normal faults in northeastern Indian Wells Valley, extends northward across eastern Coso Basin and into the southern end of Wild Horse Mesa, where it continues northward and forms a dramatic series of scarps in Pliocene volcanic flows. The faults with the largest scarps in Wild Horse Mesa exhibit a distinct left-stepping pattern (Fig. 9). A central branch, consisting of a zone of short NNE-striking, left-stepping surface traces, crosses the White Hills anticline south of Airport Lake playa and becomes the Coso Wash fault, which is characterized by a single trace along the southeastern flank of the Coso Range. A western branch of the Airport Lake fault zone crosses the southern Coso Range and joins the Little Lake fault zone in southern Rose Valley (Fig. 9). Slip Transfer across the Central Coso Range Based on the geomorphic expression of faults in late Quaternary deposits, the bulk of Holocene deformation in the Coso Range appears to be associated with the central branch of the Airport Lake fault zone (Fig. 9). The most significant structure in this branch is the Coso Wash fault, which consists of a series of NNE-striking normal faults that dip both to the SSE and the WNW. This fault zone extends from the White Hills anticline northward to Haiwee Spring in northern Coso Wash (Fig. 9) and is interpreted to be the principal locus for transferring active dextral shear through the Coso Range (Unruh et al., 2008). The Coso Wash fault zone can be traced ~9 km north of Airport Lake playa as a single SE-dipping trace that has excellent geomorphic expression as scarps in Holocene alluvial fan deposits. The fault along this reach consists of a series of alternating short NNE- and NW-striking reaches. At the southern margin of the Coso geothermal field (GF; Fig. 9), the fault splays out into a series of WNW-dipping traces that step northwest (left) into the bedrock of the Coso Range and dip toward the main geothermal production zones. The WNW-dipping fault segments are geomorphically well-expressed by NW-facing scarps in bedrock and alluvium, and the faults locally pond alluvium in their downdropped hanging wall blocks upstream of the scarps. At the latitude of the geothermal field, the E-dipping Coso Wash fault and the WNW-dipping normal faults collectively bound a prominent NNE-trending basement ridge that separates Coso Wash from the main production area of the Coso geothermal field (Walker and Whitmarsh, 1998). The basement ridge is at least 10 km long (Fig. 9), locally exhibits up to 550 m of relief,
Active tectonics of the eastern California shear zone Cenozoic
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Figure 8. Simplified geologic map of the Little Lake area (modified from Bartley et al., 2008). Distribution of granite and diorite-granodiorite units east of the Owens Valley fault zone from Whitmarsh (1998). Stereoplot shows foliation (great circles) and lineation (dots) in shear zones east of the fault zone.
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Figure 9. Northward branching of the Holocene-active Airport Lake fault zone in northern Indian Wells Valley, Rose Valley, the Coso Range, and Wild Horse Mesa. AL—Airport Lake playa; BR— basement ridge; CB—Central branch; CWF—Coso Wash fault; EB—Eastern branch; GF—geothermal field; HS— Haiwee Spring; LCF—Lower Cactus Flat; MF—McCloud Flat; UCF—Upper Centennial Flat; WB—Western branch; WHA—White Hills anticline; WHM—Wild Horse Mesa; WHMFZ— Wild Horse Mesa fault zone. Faults with especially prominent scarps in Wild Horse Mesa are highlighted in bold. Late Quaternary faults modified from Duffield and Bacon (1981) and Whitmarsh (1998), with additional original mapping. A and B indicate two faults that display evidence for late Quaternary dextral offset (please see text for discussion).
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and is best expressed between the geothermal field and Haiwee Spring. The basement ridge is essentially a horst block and may be analogous to the “central ridge” that Dooley et al. (2004) observed in scaled analog models of transtensional releasing stepovers that include a ductile substratum beneath a simulated brittle upper crust (quartz sand). North of the geothermal field, the Coso Wash fault dips consistently ESE and can be traced as a series of E-facing scarps in Holocene alluvium northward to the area around Haiwee Springs, where it loses its surface expression (Fig. 9). North of Haiwee Springs, Coso Wash terminates as a Quaternary basin and narrows to a steep canyon cut in Cretaceous bedrock and Pliocene basalts of Wild House Mesa. Analysis of stereo aerial photography of this segment of the fault indicates E-facing bedrock scarps and possibly fault-related E-facing bedrock slopes. These features probably represent Quaternary faulting, as recognized earlier by Walker and Whitmarsh (1998). The step-faulted terrain associated with the eastern branch of the Airport Lake fault and the Coso Wash fault appear to merge at this location to form a rhombic array of faults at the southern end of Upper Centennial Flat (Fig. 9), north of the termination of the basement ridge. At the latitude of Haiwee Spring, the locus of active deformation steps west from Coso Wash to a series of NNE-striking, leftstepping normal faults that bound the western margins of Quaternary basins in the northwestern Coso Range such as McCloud Flat and Lower Cactus Flat (Fig. 10). The geomorphic expression and relative activity of these structures appear to increase northward as slip dies out on the Coso Wash fault and basement ridge to the east. The left step in the locus of deformation across the Coso Range is associated with an elongated, NW-trending zone of low P-wave and S-wave velocities in the depth range of ~5–12 km (Hauksson and Unruh, 2007; Reasenberg et al., 1980; Wilson et al., 2003). The base of seismicity is distinctly elevated above the low velocity zone (Monastero and Unruh, 2002), suggesting that hot fluids (brines and/or magma) are present below ~5 km (Hauksson and Unruh, 2007; Wilson et al., 2003). At least two faults that display evidence for late Quaternary dextral offset can be traced north from Lower Cactus Flat into the northwestern Coso Range piedmont (faults A and B; Figs. 9 and 10). The easternmost of the two faults (A) transfers slip in a restraining stepover across NW-plunging, basementinvolved anticlines to the Red Ridge fault zone, a zone of short NNE-striking, en echelon normal faults that extends from Red Ridge northward to the margin of Owens Lake basin (Fig. 10). The westernmost of the two faults (B) is part of a zone of short, discontinuous NW-striking fault segments that trend toward the southern end of the Owens Valley fault (Fig. 10). Slemmons et al. (2008) documented dextral offset of Holocene beach ridges of Owens Lake along this fault trend, which they attribute to surface rupture during the 1872 Owens Valley earthquake. A Pleistocene pediment that fringes the northern Coso Range is deformed by WNW-trending folds in the triangular region between faults A and B (Fig. 10). The coeval NNE-SSW shortening and
WNW-ESE extension of the pediment surface in the triangular region between fault trends A and B is consistent with northwest dextral shear passing through the northwestern Coso Range. The Red Ridge fault zone, which is mapped in detail by Slemmons et al. (2008), may transfer slip to the Central Valley fault zone, a NW-striking fault in Owens Lake basin interpreted by Neponset Geophysical Corporation and Aquila Geosciences (1997) from analysis of seismic reflection data (Fig. 10). The “Central block” bounded by this structure and the Owens Valley fault to the west has subsided episodically during the Quaternary (Neponset Geophysical Corporation and Aquila Geosciences, 1997), possibly during large earthquakes. The NW-trending folds in the pediment surface may be the surface expression of reverse slip on blind, WNW-striking splays of the southern Owens Valley fault, also mapped by Neponset Geophysical Corporation and Aquila Geosciences (1997) from interpretation of reflection data (Fig. 10). Southern Owens Valley Fault, Northwest Coso Range The field trip stop along CA-190 (Stop 7; Fig. 1) affords an excellent opportunity to look at, stand on, and walk around evidence of surface rupture during the 1872 Owens Valley earthquake. Figure 11 is a slightly oblique aerial view of the field trip stop, looking toward the southeast. The highway parallels an abandoned shoreline of Owens Lake, and a series of gravelly late Holocene beach ridges are present on the north side of the highway. The photo shows a series of NW-trending lineaments south of the highway that project to the broad bend in the road. These features are the surface expression of the zone of discontinuous fault segments that can be traced northwest of fault A in Figure 10, and they trend toward the southern end of the Owens Valley fault zone mapped by Neponset Geophysical Corporation and Aquila Geosciences (1997) in southeastern Owens Lake basin (Fig. 10). D.B. Slemmons and his colleagues discovered that the Holocene beach ridges north of the highway are offset in a right-lateral sense along the trend of the lineaments. Slemmons et al. (2008) argue that the offset beach ridges represent surface rupture during the 1872 earthquake on the Owens Valley fault zone. If this is correct, then the 1872 rupture extended at least to the southern end of Owens Lake basin and probably south into the Coso Range piedmont (Slemmons et al., 2008). This stop is a good location to observe other tectonic-geomorphic evidence that northwest dextral shear extends from the southern end of the Owens Valley fault into the northwest Coso Range. There is a large hill ~3 km south of CA-190, on trend with the lineaments in Figure 11, where the associated fault segment terminates or makes a poorly expressed left step (Fig. 10). The hill is cut by a series of left-stepping, W-facing scarps (Fig. 12). About 3–4 km south of CA-190, there is a prominent N-facing wave-cut scarp at ~1160 m (3800 ft) elevation bordering the northern Coso Range piedmont. The shoreline of Owens Lake last reached the elevation of this escarpment ca. 24 ka during a Tioga-age pluvial highstand (Bacon et al., 2006). The shoreline
Active tectonics of the eastern California shear zone
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Structure contour (in feet) on elevation of Quaternary pediment surface (dashed where uncertain)
Figure 10. Hillshade map showing extent of a Quaternary pediment surface that fringes the northern and northwestern Coso Range. The pediment is crossed by two zones of faulting that can be traced northward from Lower Cactus flat (i.e., faults A and B). The easternmost of the fault zones (A) strikes N to NNE and joins the Red Ridge fault zone in the northern Coso Range piedmont, which in turn may merge with the “Central Valley fault zone” in southern Owens Lake basin (faults in Owens Lake basin from Neponset Geophysical Corporation and Aquila Geosciences, 1997). Fault B to the west is part of a discontinuous NW-striking zone that can be traced to the southern end of the Owens Valley fault and that experienced surface rupture during the 1872 Owens Valley earthquake (Slemmons et al., 2008). The pediment surface is folded into a series of anticlines about WNW-trending axes west of the Red Ridge fault zone.
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Shor
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Figure 11. Slightly oblique aerial view to the southeast of lineaments (arrows) associated with the southern Owens Valley fault zone (distance between arrows is ~1.1 km; see Fig. 10 for location). The field trip stop is along CA-190 at the curve in the road (Stop 7, Fig. 1). Beach ridges on the north side of the highway associated with an older Holocene shoreline of Owens Lake are offset in a right-lateral sense by the Owens Valley fault (the displacement is not visible at this scale). Slemmons et al. (2008) interpret that the displacement occurred during the 1872 Owens Valley earthquake, indicating that coseismic rupture on the Owens Valley fault extended at least as far south as the Coso Range piedmont.
Figure 12. Oblique aerial view to the east of a hill located ~3 km southeast of field trip Stop 7 on CA-190. The west flank of the hill is cut by a series of left-stepping splays of the fault shown in Figure 11. The fault splays are expressed as W-facing scarps (shadowed). Distance between the arrows is ~600 m.
Active tectonics of the eastern California shear zone scarp is visibly warped by a series of low-amplitude folds west of the Red Ridge fault zone (Fig. 10); the fold deformation is best observed from CA-190 in low-angle, early morning light when the escarpment is shadowed (Fig. 13). Although not readily visible from the field trip stop, the Red Ridge fault zone is expressed as a series of horst and graben in a middle (?) to late Pleistocene pediment surface south of the 1160 m shoreline escarpment (Figs. 13 and 14). Summary Northwest-directed dextral shear along the southeastern margin of the Sierran microplate is transferred from the Airport Lake fault to the Owens Valley fault across a discontinuous series of active structures in the central and northwestern Coso Range. Holocene surface faults in this region locally are separated by en echelon steps, and by short restraining stepovers characterized by uplift and folding. Observations by Slemmons et al. (2008) indicate that surface rupture during the 1872 earthquake on the Owens Valley fault extended at least as far south as the late Holocene shoreline of Owens Lake, and possibly southward into the Coso Range piedmont. PLEISTOCENE DEFORMATION IN NORTHERN OWENS VALLEY Several of the broad, outstanding issues concerning the pace and tempo of active deformation throughout the eastern California shear zone–Walker Lane region exist in microcosm within the
61
Owens Valley (Fig. 1). Chief among these is how one interprets differences between modern strain fields, typically measured with space geodesy (e.g., Dixon et al., 2000), and fault slip rates measured over millennia (e.g., Beanland and Clark, 1994). It has long been recognized that geodetic strain rates across the Owens Valley fault are rapid (Savage and Lisowski, 1980, 1995), and simple models of elastic strain accumulation require ~6–7 mm/yr of shear at depth to explain these observations (Gan et al., 2000). These results are in conflict with paleoseismic estimates of late Pleistocene to Holocene slip along the Owens Valley fault zone of 1–3 mm/yr (Bacon and Pezzopane, 2007; Beanland and Clark, 1994; Bierman et al., 1995; Lee et al., 2001b; Lubetkin and Clark, 1988). Recent studies explain this discrepancy as a consequence of transient post-seismic velocities induced by the 1872 Owens Valley earthquake (Dixon et al., 2003; Malservisi et al., 2001). However, long-term slip rates derived from paleoseismic data are subject to numerous epistemic uncertainties regarding the ratio of vertical to horizontal slip, uniform or non-uniform recurrence intervals, and the characteristic size of rupture events. Thus, such data are most useful when combined with geologic estimates of slip rate that average displacement over multiple seismic cycles. Active deformation across Owens Valley has also engendered considerable debate regarding longer-timescale variability of fault slip. At issue here is the question of whether coeval normal faulting along the Sierra Nevada range front and strikeslip faulting along the Owens Valley fault reflect changes in the regional stress field (Bellier and Zoback, 1995) or simply slip partitioning along a transtensional fault system (Wesnousky and Jones, 1994). Although sites where multiple, well-dated markers
INYO MOUNTAINS COSO RANGE
Figure 13. Oblique aerial view to the east of the northern Coso Range piedmont and southern Owens Valley. The darker, more eroded surface to the south (right) is a pediment surface cut across tilted strata of the Pliocene Coso Formation. The pediment terminates to the north against a wave-cut escarpment associated with one or more Pleistocene high stands of pluvial Owens Lake. The pediment and wave-cut scarp are noticeably folded about WNW-trending axes into a series of broad, low amplitude anticlines (distance between anticline axes is ~3.2 km). As demonstrated by the varying width of the shadowed escarpment, the scarp is higher across the axes of the anticlines and lower through the syncline. See Figure 10 for location of folds relative to faults cutting the pediment surface.
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Figure 14. Oblique aerial view to the southeast of the northern Coso Range pediment surface broken up in a series of horst and graben by the Red Ridge fault zone (see Fig. 10 for location). The distance along the labeled shoreline scarp is ~1 km. PLEIS TOCE NE S HORE LINE SCAR P
are displaced across a single fault are rare, several exist within the northern part of Owens Valley. Along the Fish Springs fault, a normal fault that splays from the Owens Valley fault south of the town of Big Pine (Fig. 15), throw appears to have been steady at rates of 0.2–0.3 mm/yr over the past ~300 k.y. (Martel et al., 1987; Zehfuss et al., 2001). Likewise, normal faulting along the Sierra Nevada frontal fault appears to have been steady (throw rates of 0.2–0.3 mm/yr) when averaged over the past ~120 k.y. but may have accelerated in the late Holocene (Le et al., 2007). On longer timescales, Gillespie (1991) argued that extension along the Sierra Nevada range front underwent a period of rapid slip in the middle Pleistocene, near the time of eruption of the Bishop Tuff (ca. 760 ka; Sarna-Wojcicki et al., 2000). Similarly, coordinated variations in oblique slip rate during the middle and late Pleistocene are argued to have occurred on the White Mountains fault zone and Fish Lake Valley faults (Kirby et al., 2006; Reheis and Sawyer, 1997). Resolution of the timescales over which such variations may have occurred, and the processes driving such behavior, requires increasingly precise chronology of fault slip over multiple temporal intervals. The final insight that active deformation within Owens Valley can provide with regard to the eastern California shear zone as a whole is the question of the role played by distributed arrays of faults in accommodating active deformation. Although active extension across the southern Owens Valley appears to be concentrated on the Sierra Nevada frontal fault (Le et al., 2007), numerous subsidiary faults occur throughout the northern valley (Fig. 15), suggesting the possibility that extension rates vary from south to north. Recent work in northern Owens Valley addresses aspects of each of these three issues. The key to surmounting uncertainties
in paleoseismic estimates of slip rate, and to assessing spatial and temporal variations in fault slip, is to quantify fault slip over millennial timescales, long enough to average multiple seismic cycles yet short enough to capture potential variations in fault slip. Direct dating of landscape features, enabled by cosmogenic isotopic methods (Gosse and Phillips, 2001), affords this opportunity (e.g., Frankel et al., 2007a; Kirby et al., 2006; Oskin et al., 2007). In the following, new geologic estimates of fault slip along the primary strand of the northern Owens Valley fault are presented using displaced lava flows along the eastern flank of the Crater Mountain volcanic complex (Kirby et al., 2008). Preliminary results of ongoing efforts to develop a budget of late Quaternary extension across the valley are also presented, along with a brief discussion of the regional implications of this new work. Slip Rate along the Northern Owens Valley Fault For most of its length, the trace of the 1872 rupture along the Owens Valley fault runs near or within the floodplain of the Owens River, and long-term markers of fault displacement are rare (Beanland and Clark, 1994). South of the town of Big Pine (Stop 10; Figs. 1 and 15), the fault displaces basaltic lava flows along the eastern flank of Crater Mountain (Beanland and Clark, 1994), one of the largest vent complexes in the Big Pine volcanic field. Although the similarity in composition between flows makes correlation of individual flows across the fault difficult, an apparent right-lateral separation of the contact between the flow complex and alluvial fans is present at the northeastern corner of the cone (Fig. 16). The fault geometry defines a small releasing step at this site, and the pull-apart is filled with fine-grained, young alluvium. Potential uncertainties regarding the degree of burial and
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Figure 15. Simplified tectonic map of the Owens Valley region, near the town of Big Pine (star). Map modified from Bateman (1965). Faults shown as solid lines are known to displace late Quaternary surficial deposits, whereas those shown as dashed lines are inferred to have Quaternary slip. Box shows the location of Figure 16. CM—Crater Mountain; FSF—Fish Springs fault; OVF—Owens Valley fault; RMF—Red Mountain fault.
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Figure 16. Displaced margin of the Crater Mountain flow complex near the Big Pine town dump (Stop 10, Fig. 1). (A) Geologic map of Owens Valley fault zone and Quaternary deposits overlain on air photo. Qb—Quaternary basalt; Qfo—older fan, likely equivalent to ca. 130 ka surface of Zehfuss et al. (2001); Qfy—young to active fan. (B) Retrodeformed displacement of ~235 m restores flow margin (circles). Dashed white line shows the approximate position of the buried flow margin west of the Owens Valley fault zone (after Kirby et al., 2008).
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the subsurface geometry of the flow margin required assessment before this site could provide a robust constraint on fault slip. To test the hypothesis that dextral separation of the flow margin observed at the surface (235 ± 15 m) is a reliable estimate of lateral slip along this segment of the Owens Valley fault, a ground-penetrating radar survey of the site was conducted (Kirby et al., 2008). The results of this survey confirm the presence of a shallowly buried flow margin to the west of the fault (Fig. 16), but do not reveal any basalt beneath the pull-apart itself. Rather, the flow margin appears to end abruptly against alluvium. Thus, Kirby et al. (2008) concluded that the horizontal component of fault slip at this site is ~235 m. The age of the flow surface was estimated using the concentrations of cosmogenic 36Cl in samples taken from well-preserved remnants of the flow surface. Samples were selected to minimize the chance of burial by eolian or alluvial material and were collected from outcrops that exhibited glassy surfaces representative of minimal surface lowering during weathering. Three samples were collected locally, on the west side of the fault, and three additional samples were collected from flows on the southwestern side of the vent complex, near the Red Mountain fault (Fig. 15). Because perfect surface preservation in this environment is extremely unlikely, exposure ages were modeled for a range of possible lowering rates (Kirby et al., 2008), and the resultant age ranges are taken as a best estimate of the age of the flow. Ages from both sample localities overlap within uncertainty, and indicate that the flow is 70 ± 14 ka (Kirby et al., 2008). When combined with the displacement of the flow margin, these results imply that slip rates on the Owens Valley fault during this time period have been 3.6 ± 1.0 mm/yr. Notably, this range is consistent with, but at the high end of, previous estimates (Beanland and Clark, 1994; Lee et al., 2001b), and it is 2–3 times greater than Holocene estimates of Bacon and Pezzopane (2007). Whether these differences reflect spatial differences in slip between the northern and southern segments of the fault zone, or whether they represent a period of rapid slip prior to ca. 20–25 ka remains unknown (Kirby et al., 2008). Distributed Extension across Northern Owens Valley Ongoing work is focused on developing a budget for late Pleistocene extension across the northern part of Owens Valley. Here, a limited portion of this effort is discussed, focused on new estimates of the slip rate along the Red Mountain and Birch Mountain faults (Fig. 15). The Red Mountain fault is a 10-km-long normal fault that extends south from the southwestern flank of Crater Mountain, subparallel to and ~2 km west of the Fish Springs fault (Stop 9; Figs. 1 and 15). It is marked along much of its trace by W-facing scarps that pond young alluvium in the hanging wall block. Despite its length and proximity to the Fish Springs fault, the Red Mountain structure has received considerably less attention, and the role of this fault in the extension budget across northern Owens Valley is essentially unknown.
At the northern end of the fault, lava flows from Crater Mountain are displaced in a W-side–down sense and have been subsequently buried by young alluvial fans emanating from Birch Creek. The flow itself buries an older alluvial fan surface with a moderately well-developed soil profile; surface characteristics of this fan suggest that it is likely equivalent to surfaces dated 2 km to the east at ca. 135 ka (Zehfuss et al., 2001), whereas the younger fan appears continuous with surfaces dated at 13–15 ka (Zehfuss et al., 2001). Chlorine-36 ages from the footwall exposures of the flow cluster at 70 ± 14 ka (Kirby et al., 2008) and are consistent with the regional chronology of Zehfuss et al. (2001). The flow surface has been offset vertically by 14.5 ± 1.5 m, and two small scarps are present in the youngest late Pleistocene surface that exhibit a combined throw of 2.5 ± 0.2 m. Both of these displacements are consistent with long-term vertical displacement rates along the Red Mountain fault of 0.2–0.3 mm/yr (Greene et al., 2007). Thus, this structure appears to play as great a role in the regional deformation field, as does the Fish Springs fault. Significant slip is also present along the Sierra Nevada frontal fault at this latitude, as indicated by prominent scarps high on the range front (Fig. 17). The fault displaces sharp, triangular moraine crests in both the Tinemaha and Birch Creek drainages; moraines appear fresh, with minimal soil development and little to no weathering of boulder surfaces. Vertical displacements in both drainages are similar, ranging from 7 to 9 m (Fig. 17). Exposure ages derived from cosmogenic 36Cl concentrations in five samples taken from boulders atop the northern lateral moraine in Tinemaha Creek cluster between 14–15 ka (Greene et al., 2007). Thus, vertical displacement rates on this segment of the Sierra Nevada frontal fault system appear to be relatively high, at 0.5–0.7 mm/yr. These rates may reflect relatively recent breaching of the relay between the northern tip of the southern Sierra Nevada frontal fault system and the southern tip of the Round Valley fault. Regional Implications From a regional perspective, these results contribute to an emerging picture of spatial variations in slip rate along the eastern California shear zone. Recent estimates of late Pleistocene slip rates along the Death Valley–Fish Lake Valley fault system suggest that slip rates decrease northward, from 4 to 5 mm/yr along the central section (Frankel et al., 2007a) to 2–3 mm/yr along the northern sections (Frankel et al., 2007b). In a similar fashion, right-lateral slip rates in the Owens Valley appear to decrease northward, from 3 to 4 mm/yr along the Owens Valley fault to ~0.5 mm/yr along the White Mountain fault zone (Kirby et al., 2006). These apparently systematic patterns suggest that simple approaches to reconciling geologic slip and geodetic strain by summing geologic slip along a two-dimensional transect are likely to be misleading. Although strain accumulation and release may reconcile across a given transect (e.g., Frankel et al., 2007a), the spatial distribution of strain in this young, evolving fault system appears be quite variable along strike.
Active tectonics of the eastern California shear zone
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Figure 17. Topographic surveys of displaced moraine crests along the Birch Mountain segment of the Sierra Nevada frontal fault zone. Surveys were conducted using post-processed differential global positioning system (sub-cm precision). Cosmogenic 36Cl ages from boulders on the Tinemaha Creek moraine indicate these deposits are 13–15 ka (Greene et al., 2007).
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Summary New estimates of late Quaternary deformation rates in the vicinity of Big Pine, California suggest that (1) slip rates along the northern Owens Valley fault zone over the past 55–85 k.y. are 3.6 ± 1.0 mm/yr, significantly higher than recent estimates along the southern portion of the fault zone, but similar to geodetic data, and (2) that extension rates associated with distributed fault arrays in the northern Owens Valley also appear to be higher than recent estimates in the southern Owens Valley. These results highlight spatial variations in late Quaternary slip rates throughout the eastern California shear zone and suggest caution in comparison of geologic slip rates with geodetic data along 2-D transects. SPATIAL VARIATIONS IN LATE PLEISTOCENE SLIP RATE ALONG THE DEATH VALLEY–FISH LAKE VALLEY FAULT ZONE The Death Valley–Fish Lake Valley fault zone is the largest and most continuous fault system in the eastern California shear zone, extending some 300 km northward from its intersection with the Garlock fault (Fig. 1). Both geologic and geodetic observations suggest that the Death Valley–Fish Lake Valley fault zone accommodates the majority of slip in the northern eastern California shear zone. Specifically, several space-based geodetic surveys show that, over the past ~15 yr, the Death Valley–Fish Lake Valley fault zone has been taking up 3–8 mm/yr of the measured 9.3 ± 0.2 mm/yr of Pacific–North America plate motion in the northern eastern California shear zone and Walker Lane (Bennett et al., 1997, 2003; Dixon et al., 1995, 2000, 2003; Humphreys and Weldon, 1994; McClusky et al., 2001; Savage et al., 1990; Wernicke et al., 2004). Along the northern part of the Death Valley–Fish
Lake Valley fault zone in Fish Lake Valley, modeling of geodetic data shows the fault system is storing strain at 4–10 mm/yr and that the White Mountains fault zone, the other major strike-slip structure at this latitude, stores strain at 1–5 mm/yr (Dixon et al., 1995, 2000). The geodetic data, therefore, suggest that at latitude ~37.5°N, essentially all plate boundary strain in the eastern California shear zone is accommodated on these two structures. Although numerous geodetic data are available from the region, only a few field-based studies have attempted to measure intermediate- and long-term (1,000–1,000,000 yr) geologic slip rates on the Death Valley–Fish Lake Valley fault zone (Brogan et al., 1991; Frankel, 2007; Frankel et al., 2007a, 2007b; Klinger, 2001; Reheis and Sawyer, 1997). A lack of geochronologic constraints on the age of offset alluvial landforms has thus made it difficult to compare rates of deformation over multiple time scales along this part of the plate boundary. Previous estimates of the late Pleistocene slip rate for the Death Valley–Fish Lake Valley fault zone in Fish Lake Valley range from 1 to 9 mm/yr (Reheis and Dixon, 1996; Reheis and Sawyer, 1997). However, recent work combining high-resolution ALSM digital topographic data with cosmogenic nuclide geochronology to investigate fault offsets and slip rates has helped refine the long-term slip rates in this region (Frankel, 2007; Frankel et al., 2007a, 2007b). These new results reveal spatial variations in the slip rate along the Death Valley–Fish Lake Valley fault zone and have important implications for eastern California shear zone and Pacific–North America plate boundary kinematics. Faulting in Fish Lake Valley The Death Valley–Fish Lake Valley fault zone bounds the east side of the White Mountains in Fish Lake Valley (Fig. 1).
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This region makes up the northern 80 km of the Death Valley–Fish Lake Valley fault zone. Estimates of total dextral displacement along the northern Death Valley–Fish Lake Valley fault zone are thought to range from ~50 to 80 km since Cambrian to Middle Jurassic time (McKee, 1968; Stewart, 1967). More recent right lateral–oblique fault activity in Fish Lake Valley is characterized by numerous deformed geomorphic features, including fault scarps, displaced alluvial fans, offset drainage channels, shutter-ridges, and sag ponds (e.g., Brogan et al., 1991). An ALSM survey was recently conducted along the Death Valley–Fish Lake Valley fault zone in northern Death Valley and Fish Lake Valley to precisely map displacement on late Pleistocene and Holocene fault-related landforms (please see Frankel, 2007, and Frankel et al., 2007b for details about ALSM data collection and processing parameters). Results from two of the survey locations, the offset Furnace Creek and Indian Creek alluvial fans in Fish Lake Valley, are discussed below.
Furnace Creek Offset The Furnace Creek alluvial fan in central Fish Lake Valley is located along the Oasis section of the Death Valley–Fish Lake Valley fault zone (Stop 11, Fig. 1; Brogan et al., 1991; Reheis and Sawyer, 1997). At this location, alluvial fans are offset along two parallel strands of the fault (Fig. 18). Previous work established a number of possible late Pleistocene channel displacements at Furnace Creek, ranging from 111 to >550 m (Brogan et al., 1991; Reheis et al., 1995; Reheis and Sawyer, 1997). ALSM data were used to determine more precise offset measurements. These high-resolution topographic data helped reveal subtle topographic features with hillshade, topographic, and slope aspect maps, topographic profiles, and thalweg positions to reconstruct a prominent drainage channel and the overall fan morphology (Fig. 18). Based on these data, Frankel et al. (2007b) revised the late Pleistocene strike-parallel displacement for the Furnace Creek alluvial fan to 290 ± 20 m (Fig. 18).
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Figure 18. Maps of Furnace Creek alluvial fan in Fish Lake Valley (Stop 11, Fig. 1; after Frankel et al., 2007b). (A) Hillshaded 1-mresolution ALSM digital elevation model of the Furnace Creek alluvial fan. Star indicates parking area for field trip. (B) Geologic map (modified from Reheis et al., 1995) of the Furnace Creek alluvial fan draped over the hillshaded image from A. Qfio—Older alluvium of Indian Creek (late Pleistocene); Qft—Alluvium of Trail Canyon (middle Pleistocene); Qfm—Alluvium of McAfee Creek (middle Pleistocene). (C) Furnace Creek alluvial fan retro-deformed 290 ± 20 m to its prefaulting position based on the high-resolution ALSM digital elevation data. Hatched pattern on northwest section of the offset fan indicates a surface of similar age, but set into the Qfio unit. Combining the late Pleistocene displacement history with the 94 ± 11 ka cosmogenic 10 Be model age from the offset Qfio surface yields a slip rate of 3.1 ± 0.4 mm/yr (Frankel et al., 2007b).
Active tectonics of the eastern California shear zone Indian Creek Offset The Indian Creek alluvial fan is located at the north end of Fish Lake Valley, near the northern termination of the Death Valley–Fish Lake Valley fault zone (Stop 12, Fig. 1). The fan is part of the Chiatovich Creek section of the Death Valley–Fish Lake Valley fault zone (Brogan et al., 1991; Reheis and Sawyer, 1997). Here, the fault zone splays into numerous normal faults, and the strike-slip component of deformation is localized along a single strand that displaces both late Pleistocene and Holocene alluvium (Fig. 19). Reheis et al. (1993) and Reheis and Sawyer (1997) estimated 83–165 m of late Pleistocene rightlateral deformation at this location based on a single offset debris flow channel. Hillshade, slope aspect, and topographic maps and channel thalwegs derived from ALSM data were used to revise the late Pleistocene displacement at Indian Creek to 178 ± 20 m by retrodeforming at least four, and possibly six, offset channels incised through the alluvial fan (Fig. 19; Frankel et al., 2007b).
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Alluvial Fan Ages Ages of the offset alluvial fans in Fish Lake Valley were previously estimated on the basis of soil development and surface morphology (e.g., Reheis et al., 1993, 1995; Reheis and Sawyer, 1997). Both the Furnace Creek and Indian Creek fans have similar soil and morphologic characteristics. The offset late Pleistocene fans have well-developed soils with a 5–10-cm-thick silty vesicular A horizon and an argillic B horizon with moderate clayfilms and stage II to III carbonate development (Reheis and Sawyer, 1997). The fan surfaces are characterized by a moderately to well-developed desert pavement, moderate to dark desert varnish coatings on clasts ranging in size from pebbles to boulders, and subdued to moderately incised channels (Fig. 20). Ages of the offset alluvial fans at Furnace Creek and Indian Creek were quantified by cosmogenic nuclide 10Be geochronology (Frankel et al., 2007b; Gosse and Phillips, 2001). Geochronology samples were collected from the top 2–5 cm of large
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Figure 19. Maps of the Indian Creek alluvial fan in Fish Lake Valley (Stop 12, Fig. 1; after Frankel et al., 2007b). (A) Hillshaded 1 m resolution ALSM digital elevation model of the Indian Creek alluvial fan. Star indicates parking area for field trip. (B) Geologic map (modified from Reheis et al., 1993) of the Indian Creek alluvial fan and surrounding areas draped over the hillshaded image in A. Qfiy—Younger alluvium of Indian Creek (late Pleistocene); Qfl—Alluvium of Leidy Creek (early Holocene and late Pleistocene); Qls—Holocene to late Pleistocene landslide deposits. (C) Indian Creek alluvial fan retrodeformed 178 ± 20 m to its pre-faulting position based on ALSM data. A cosmogenic 10Be model age for the Qfiy surface of 71 ± 8 ka yields a slip rate of 2.5 ± 0.4 mm/yr when combined with the late Pleistocene offset measurement (Frankel et al., 2007b).
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Figure 20. Photograph looking west across the Furnace Creek alluvial fan. The White Mountains make up the distant skyline ridge. Boulder in the foreground is representative of locations from which samples were collected on the Furnace Creek and Indian Creek alluvial fans. Note the high degree of varnish development on the boulder surface. Boulders averaged ~115 × 90 × 85 cm at Furnace Creek and ~65 × 55 × 50 cm at Indian Creek.
boulders on the stable parts of offset fan surfaces mapped as Qfi by Reheis et al. (1993, 1995; Fig. 20) and analyzed by accelerator mass spectrometry at Lawrence Livermore National Laboratory (please see Frankel, 2007, and Frankel et al., 2007b, for geochronology methods and data). Furnace Creek Fan Age Eight cosmogenic 10Be samples from the Qfio surface at Furnace Creek (Fig. 18; unit Qfi of Reheis et al., 1995) range in age from 79 ± 8 ka to 112 ± 8 ka (Frankel et al., 2007b). The relatively tight cluster of dates suggests that the Qfio surface has remained stable, with samples having been exposed to cosmic rays in their current configuration, since deposition. The age of the fan is taken to be the mean and standard deviation of the eight samples, which yields a date of 94 ± 11 ka (Frankel et al., 2007b). This age falls within the broad age range of 50–130 ka estimated for the Furnace Creek fan on the basis of soil development and surface morphology by Reheis and Sawyer (1997). Indian Creek Fan Age Eight cosmogenic 10Be samples were also collected from the Qfiy surface (Fig. 19; unit Qfi of Reheis et al., 1993) at Indian Creek. These samples range in age from 59 ± 4 ka to 81 ± 6 ka (Frankel et al., 2007b). The samples from Indian Creek are somewhat younger than those from Furnace Creek, yet they display a similarly tight cluster, which is also interpreted to indicate the stability of the Qfiy surface since deposition. The mean and standard deviation of the eight 10Be dates yields an age of 71 ± 8 ka for the Qfiy surface at Indian Creek, which is in agreement with the previously estimated range of 50–130 ka (Frankel et al., 2007b; Reheis and Sawyer, 1997).
The cosmogenic 10Be surface exposure dates from the Furnace Creek and Indian Creek fans are interpreted as maximum ages for the deposits because the channels used as piercing points must have formed at some undetermined time following fan deposition. The slip rates reported here should therefore be interpreted as minima, though it appears that most, if not all, of the rightlateral deformation accommodated on faults has been accounted for. The pervasive normal faulting observed on the eastern White Mountains piedmont is not considered in these rates. Combining the displacement of 290 ± 20 m with the age of 94 ± 11 ka for the offset Qfio surface at Furnace Creek yields a late Pleistocene slip rate for the Oasis section of the Death Valley–Fish Lake Valley fault zone of 3.1 ± 0.4 mm/yr. Previous slip rate estimates at this location ranged from 1.5 to 9.3 mm/yr (Reheis and Sawyer, 1997). The 178 ± 20 m of displacement and 71 ± 8 ka age of the offset Qfiy surface at Indian Creek results in a slightly slower slip rate of 2.5 ± 0.4 mm/yr along the northern Chiatovich Creek section of the Death Valley–Fish Lake Valley fault zone. This rate falls within the bounds of the previous slip rate estimate of 1.1–3.3 mm/yr for this site (Reheis and Sawyer, 1997). Spatial Variations in Slip Rate and Northern Eastern California Shear Zone Strain Distribution Previous studies in the northern eastern California shear zone suggest that the down-to-the-NW faults located between the major strike-slip faults of the region transfer slip from the Owens Valley and Panamint Valley–Hunter Mountain–Saline Valley faults to the northern part of the Death Valley–Fish Lake Valley fault zone (Fig. 1; Dixon et al., 1995; Lee et al., 2001a; Reheis and Dixon, 1996). However, recent results from three slip rate sites along the Death Valley–Fish Lake Valley fault zone show that this may not be the case (Frankel et al., 2007a, 2007b). The late Pleistocene slip rate along the Death Valley–Fish Lake Valley fault zone in northern Death Valley is ~4.5 mm/yr (Frankel et al., 2007a). The rates determined at the offset Furnace Creek and Indian Creek alluvial fans show that this rate decreases to ~2.5–3 mm/yr on the northern part of the Death Valley–Fish Lake Valley fault zone in Fish Lake Valley (Frankel et al., 2007b). The late Pleistocene slip rate on the White Mountains fault zone is 0.3–0.4 mm/yr (Kirby et al., 2006). Taken together, the late Pleistocene slip rates on the two major faults at latitude ~37.5°N are less than half the 9.3 ± 0.2 mm/yr region-wide rate of dextral shear determined from geodetic data (Bennett et al., 2003). This result suggests either that deformation at the latitude of Fish Lake Valley is accommodated on structures other than the White Mountains fault zone and Death Valley–Fish Lake Valley fault zone or that a strain transient exists in the northern eastern California shear zone, similar to that proposed for the Mojave Desert (Oskin and Iriondo, 2004; Oskin et al., 2006, 2007). Strain rates appear to have remained constant in the northern eastern California shear zone over the past ~70 k.y. at the
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latitude of northern Death Valley (Frankel et al., 2007a). If true, this implies that strain must be accommodated off the White Mountains and Death Valley–Fish Lake Valley fault zones, either to the west through Long Valley caldera (e.g., Kirby et al., 2006) or to the east from the Emigrant Peak fault zone through the Silver Peak–Lone Mountain extensional complex (Fig. 1; Oldow et al., 1994). If the faults to the east are indeed accommodating this additional strain, then the eastern California shear zone–Walker Lane transition must occur farther south, in a broader, more diffuse zone than previously recognized.
FIELD GUIDE
Summary
Day 1: Mojave Desert, Summit Range, and Red Rock Canyon
The late Pleistocene slip rate on the Death Valley–Fish Lake Valley fault zone decreases from ~4.5 mm/yr in northern Death Valley to ~3 mm/yr at Furnace Creek in central Fish Lake Valley and ~2.5 mm/yr at Indian Creek in northern Fish Lake Valley. Combining slip rates from the northern Death Valley–Fish Lake Valley and White Mountains fault zones, the two major faults at latitude ~37.5°N, indicates that the late Pleistocene rate of dextral shear is less than half that determined from geodetic data. This suggests either the existence of a strain transient in the northern eastern California shear zone or that deformation is distributed across other structures in the region.
Stop 1: Lenwood Fault (Southern Offset) Directions. This stop is located at UTM 507142E, 3844558N in Stoddard Valley, east of highway 247, ~10 mi south of Barstow, California. To reach the site from Interstate 15, exit at CA247 (Barstow Road). Turn south on CA-247 and drive ~10 mi. The field site is located on Bureau of Land Management (BLM) route OM8, which is ~1 mi south of the Slash X Ranch Café, on the left. To locate the road, look for the end of barbed wire fence and a “Call Box” sign. Turn left on to BLM route OM8 immediately after the “Call Box” sign. There are several homesteads on OM8 on the way to the site, so please drive slowly to keep the dust down and watch for oncoming traffic. Take OM8 for ~4.5 mi. You will pass over a cattle guard. If you drive through a creek bed, then up a steep rocky section of the road ~4 m high, you’ve probably just passed it. Park in the flat area just before this rocky slope (Fig. 2). Description. Figure 2 shows a hillshade map of this stop generated from a portion of the Lenwood fault ALSM survey (also see Fig. 21). A feature of particular interest at this location is a 100-m-wide pull-apart basin with slip partitioned onto several strands. A channel that crosses this basin appears offset ~45 m. Northwest of the pull-apart basin, fault displacement
SUMMARY The eastern California shear zone accommodates the majority of Pacific–North America plate boundary motion east of the San Andreas fault. Therefore, deformation in this region is critical to our understanding of plate boundary kinematics, in addition to the behavior and evolution of fault systems. The studies presented in this guidebook are some of the most recent investigations undertaken in the region. We have highlighted new late Pleistocene slip rates in Owens Valley, Fish Lake Valley, and the Mojave Desert; discussed the timing and rates of offset along the Garlock fault; provided evidence for ~65 km of total cumulative right-lateral displacement across Owens Valley; and observed recent deformation associated with southern Owens Valley fault in the northwestern Coso Range. Together, the topics in this guidebook span the Cretaceous to late Holocene tectonic history of the eastern California shear zone. As is often the case, much of this recent research, while generating a wealth of new data on slip rates, displacement histories, and fault kinematics, has brought about even more questions regarding spatial and temporal patterns of deformation in the eastern California shear zone. The eastern California shear zone provides a unique opportunity to study the kinematics of an evolving plate boundary. Continued work in the region, focused on defining rates of deformation across broad temporal (tens to millions of years) and spatial (tens to hundreds of kilometers) scales, will help fill gaps in our understanding of the role the eastern California shear zone plays in accommodating Pacific–North America plate boundary deformation.
This guide provides directions to field trip stops relative to nearby towns, landmarks, and major roads. It does not provide distances between individual stops. All field trip stop locations are keyed to the map in Figure 1. In addition, Universal Transverse Mercator (UTM) coordinates of all stops are provided in NAD83 datum, zone 11. It is preferable to have a high-clearance vehicle (and possibly 4WD, depending on road conditions) for many of the stops.
Figure 21. Low-relief scarps of the Lenwood fault (in mid-ground; Stop 1). Photograph taken looking to the northeast.
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formed a prominent SW-facing scarp cutting both Q2a and Q2b alluvial fans. Inset into the fans are younger channel-fill deposits also cut by low, SW-facing scarps in several localities. One of the larger channels crossing the fault is deflected ~270 m. Most of this deflection is due to construction of a Q2b alluvial fan in front of the channel. Two secondary faults form scarps in the late Pleistocene fans east of the main fault trace.
valley. The slip rate of the Lenwood fault was estimated here from deflection of a pair of channels located immediately NW of the road, and from a shutter ridge forming in front of the larger channel descending the fault-line valley. These channels are inset into outcrops of conglomerate and Q2b alluvial fan surfaces. Based on the 37 ± 7 ka age of Q2b determined from the southern site, the slip rate here is 0.8 ± 0.2 mm/yr.
Stop 2: Lenwood Fault (Northern Offset) Directions. This stop is located at UTM 501464E, 3851618N. The stop is ~3.5 mi from the highway. From Stop 1, return to CA-247 and turn right to go north. Approximately 1.5 mi north of the Slash X Ranch Café, bear right onto Stoddard Valley Road. The intersection is located 0.2 mi past the power lines with towers shaped like a Π symbol. Take this road straight back with no turns until you are past the first set of hills. Here, look for a dirt track cutting southeast, up the valley, to join the power-line road. Turn left onto the power-line road and ascend the hills. Note that it is also possible to drive the power-line road from CA-247; however, this route requires a very steep descent of the first set of hills. Once on the power-line road, stay left at the next three forks. Park at the fourth and final fork, located at the crest of the hills (Fig. 3). The Lenwood fault runs along the hillside, ~80 m below. Walk down the road to reach the fault. It is not recommended to drive down the hill, as the descent is dangerously steep and may require 4WD to ascend. Description. Stop 2 features an ~600 m length section where the Lenwood fault is spectacularly exposed as a set of uphill-facing scarps and shutter ridges (Figs. 3 and 22). Southward, the fault crosses a wash and runs up the northeast side of a fault-line
Stop 3: Garlock Fault in Summit Range Directions. This stop is located at UTM 445564E, 3924572N. This location is between Barstow and Ridgecrest, near the small town of Johannesburg. From Barstow, head east on CA-58 until it intersects with US-395. Go north on US-395 for ~27 mi to Trona Road. Turn east on Trona Road and continue for ~7.75 mi to the dirt road leading off to the west; park there. Description. This stop is to introduce the main units in the Summit Range. This will be the basis for correlation/comparison with rocks in the Red Rock Canyon area (Stop 4). We will walk through the stratigraphy here to examine the various units (Figs. 5 and 23). First stops will be in the lower sedimentary sequence. We will then examine the lapilli tuff and its relation to underlying and overlying units. Lastly, we will look at the various tuffs overlying the lapilli tuff. Stop 4: Garlock Fault in Red Rock Canyon Directions. This stop is located at UTM 411017E, 3913555N. Red Rock Canyon is located 25 mi north of Mojave on CA-14, near Cantil. Signs on CA-14 clearly indicate the turnoff on Abbott Road. Go west on Abbott Road to the visitor center parking lot and park. From there, walk ~0.2 mi northeast to the outcrop.
Figure 22. View looking to the northwest at fault scarps and shutter-ridges along the Lenwood fault (Stop 2). Note truck on right side of photograph for scale.
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Description. This stop will emphasize units of the Dove Spring Formation (Figs. 5 and 24). We start at the parking area east of the road and proceed up section. Our first stops are in the sedimentary and volcanic units below the lapilli tuff. First outcrops are the conglomerates in the Td2 sequence. We will then work our way up-section into white tuffs below the thick lapilli tuff. Proceed then through the tuff looking at the possible cooling break in the middle of this massive unit. We will then look at (or hike westward to) the rocks above the tuff, including the basalt flows in the higher parts of the section. Day 2: Little Lake, Cactus Flat, Coso Range, Independence Creek, and Birch Creek Stop 5: Independence Dike Swarm at Little Lake Directions. This stop is located at UTM 417369E, 3977338N, just west of Little Lake. Heading north on US-395, exit at the southernmost Little Lake Road exit. (NOTE: if approaching this stop from the north, this will be the second Little Lake Road exit; if approaching from the south, this will be the first Little Lake Road exit.) Cross to the west side of the highway and drive 0.1 mi for a view stop. Description. This is stop 13 of Glazner et al. (2005). The Little Lake structural block (Fig. 8) exposed at this stop and in the mountains directly to the east is composed of Jurassic (165 Ma; Whitmarsh, 1998) plutonic rocks that vary in composition from diorite to leucogranite (mainly as a result of pervasive magma mingling) and lack a penetrative fabric. The heterogeneous plutonic complex is intruded by Jurassic (148 Ma; Whitmarsh, 1998) Independence dikes that vary in composition roughly as much as their wall rocks do. Because the dikes are more erosion-resistant than their wall rocks, most of the craggy outcrops on the hills to the east are dikes. A traverse of the skyline to the southeast of the stop yielded 9.1% dilation by dike intrusion (Fig. 8; Bartley et al., 2008). Individual dikes in the Little Lake block range up to >10 m thick, yet any individual dike cannot be traced for more than
Figure 23. Photograph looking southeast across the Summit Diggings volcanic field (Stop 3). The prominent ridge in the center of the photo is a N80E-striking, SSE-dipping sequence of tuffs capped by a distinctive pink lapilli ash-flow tuff. This outcrop is surrounded by a patchwork of tuffs, epiclastic rocks, and sedimentary rocks of the Bedrock Springs Formation. The photo is taken from the side of a dacite dome (dark-colored rocks in left foreground) which is part of a larger 12– 11 Ma dome-flow-tuff complex. The tuffs occur in a topographically low valley that is likely related to formation of a small caldera resulting from the extrusion of the pink ash-flow tuff.
200 m because the dikes are transected by numerous NW- to N-striking ductile-brittle shear zones. The complete mismatch of dikes across each shear zone suggests that displacements across individual zones may be large. Shear zones commonly are exposed in numerous prospect pits that dot the area, and consist of greenschist-facies (white mica–chlorite–albite–epidote ± biotite) phyllonite. Most of the shear zones dip steeply to moderately westward (Fig. 9), and lineation and shear-sense indicators (S-C composite foliations, asymmetric porphyroclasts, mica fish)
Figure 24. An ~125 m sequence of tuffs, epiclastic, and clastic units of the lower Dove Spring Formation, Ricardo Group, located on the north side of the Garlock Fault in the Red Rock Canyon State Park (Stop 4). The E-dipping normal fault through the sequence displaces the units ~25 m. The prominent white air fall tuffs are dated at older than 11 Ma. The capping pink lapilli ash-flow tuff correlates with the pink tuff found at the Summit Range, 34 km to the east on the south side of the Garlock Fault. The red and tan clastic units are also identical in lithology to those found in the Summit Range area that are attributed to the Bedrock Springs Formation.
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indicate varying proportions of dextral and reverse displacement (Bartley et al., 2008). Shear zones decrease in abundance eastward toward the main axis of the Coso Range. Timing relations, kinematics, and spatial distribution of the shear zones suggest that they may be representative of structures that accommodated significant offset across Owens Valley. Roadcuts along the microwave tower access road south of this stop expose rocks of the Sierran block, including granodiorite and granodiorite gneiss, variably intruded by 10–20-cm-thick mafic dikes, all of which are thoroughly shattered. Like plutonic rocks of the Little Lake block, the granodiorite yielded a U-Pb age near 165 Ma (Bartley et al., 2008), but geologic similarities between the Little Lake block and adjacent Sierra rocks end there. The granodiorite varies much less in composition than Little Lake plutonic rocks, contains a pervasive gneissic fabric that is locally migmatitic, and contains quartzite and calc-silicate rock xenoliths not seen in the Coso Range. Dikes in the granodiorite are uniformly mafic, rarely as thick as 1 m, vary widely in orientation, and account for only 1%–2% dilation. In areas where geochronology and cross-cutting relations distinguish Jurassic Independence dikes from associated Cretaceous dikes (Coleman et al., 2000; Mahan et al., 2003), all of these characteristics correspond to Cretaceous dikes. Note that rocks of the Little Lake block are exposed on the west side of US-395 in low outcrops along the west side of Little Lake Road immediately north of here (on and around hill at UTM 417979E, 3977298N). There is an almost complete geologic mismatch between the Little Lake block and the adjacent Sierra Nevada, and it is inferred that the boundary in between, which is concealed by alluvium, is a locus of major tectonic offset.
Stop 6: Coso Dike at Cactus Flat Directions. This stop is located at UTM 418900E, 4008439N, ~5 mi southeast of Olancha. From Olancha, drive south on US395 ~1.5 mi to Cactus Flat Road (opposite the fire station). Turn east and drive past the pumice mine (mine is 4.2 mi from intersection of Cactus Flat Road and US-395). Two and a half miles beyond the mine, make a left turn on a narrow dirt road, and 0.4 mi beyond this, turn left at a side road at UTM 419210E, 4007018N. Continue down this road ~0.5 mi and park at UTM 419490E, 4007610N. An outcrop of the Coso dike is obvious in the hillside north-northwest of the parking area. Description. This is Stop 11 of Glazner et al. (2005). Cretaceous dikes in the Coso Range were first described by Duffield et al. (1980) as steeply dipping, W-striking, conspicuously K-feldspar and quartz-phyric granitic dikes that are up to 10 m thick. Whitmarsh (1998) named these dikes the Coso dike swarm, mapped the dikes in significant detail, and obtained a preliminary U-Pb zircon date of ca. 84 Ma from a thick dike in the swarm exposed near Upper Centennial Flat. One of us (Glazner) made a reconnaissance trip into the Coso Range in 1996 and noted the strong similarity between the Coso and Golden Bear dikes and proposed that the two areas might once have been contiguous. Kylander-Clark (2003) tested this correlation and found identical ages of 83.5 Ma for the two swarms, overlapping geochemistry, and similar, unique wall rocks (102 Ma leucogranite) that strongly support correlation of the two areas and ~65 km of dextral displacement. At this locality, the Coso dike (Fig. 25) is >20 m wide, strikes 281°, and dips ~80°N. The quartz monzonite porphyry contains 1–5-cm-long euhedral K-feldspar phenocrysts that locally contain Carlsbad twins. Plagioclase occurs as small (2– 4 mm) subhedral, subequant grains. Quartz occurs in euhedral
Figure 25. Westernmost outcrop of Coso dike swarm in the Coso Range (Stop 6). View is to the northwest from Cactus Flat. The dike is ~20 m thick, strikes W, and forms the prominent crags on skyline; the skyline behind is underlain by Pliocene mafic volcanic rocks that overlap the dike.
Active tectonics of the eastern California shear zone bipyramids that range from 2 to 8 mm in diameter. Biotite is the major mafic mineral, occurring as small (typically 1 mm) subhedral to euhedral grains. Looking east, the dike continues on strike, but is difficult to follow from a distance. Stop 7: Northern Coso Range Piedmont Directions. This stop is located at UTM 413630E, 4019852N. From Olancha, head east on CA-190, ~3.5 mi from the intersection with US-395 (~1 mi west of the well-marked road to Dirty Sock Corporation Yard) and park on the side of the road along the broad shoulder. Description. The thing to look at, stand on, and walk around here is evidence for surface rupture during the 1872 Owens Valley earthquake. With a little searching, you will find a series of rightlaterally displaced Holocene beach ridges on the north side of the highway that were discovered by Burt Slemmons and his colleagues ~30 years ago and that are on trend with the Owens Valley fault to the north where it heads into the playa from Bartlett Point (Fig. 11). Slemmons et al. (2008) argue that the offset beach ridges indicate that surface rupture during the 1872 earthquake extended all the way to the Coso Range piedmont, rather than dying out at Bartlett Point as suggested by Beanland and Clark (1994). Once you find the offset beach ridges, look to the southeast to view tectonic-geomorphic evidence for dextral shear extending from the southern end of the Owens Valley fault into the northwest Coso Range. There is a large push-up ridge ~3 km south of the road where the fault zone that offsets the beach ridges terminates or makes a left step. About 3–4 km south of the road is a high, N-facing wave-cut scarp at ~1160 m (3800 ft) elevation along the northern Coso Range piedmont. This scarp was last inundated ca. 24 ka during a Tioga-age pluvial highstand of
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Owens Lake. The wave-cut scarp is visibly warped into a series of low-amplitude folds associated with a fault zone that transfers slip from Owens Valley southward through the Coso Range to the Airport Lake fault zone in Indian Wells Valley. These structures comprise the eastern tectonic margin of the Sierra Nevada microplate at this latitude. Stop 8: Golden Bear Dike Near Independence Creek Directions. This stop is located at UTM 387410E, 4068329N along the Sierra Nevada range front southeast of Independence. From the center of Independence, drive 4.5 mi west on Onion Valley Road. Turn south on Foothill Road and continue for 1.5 mi to the foot of the boulder-covered hill and park on the side of the road. Description. This is Stop 6 of Glazner et al. (2005). The Golden Bear dike crops out from the valley floor to the range crest and beyond, crossing the crest north of Forester Pass to eventually peter out in the headwaters of the Kern River (Fig. 6). Looking west from this locality, it is possible to trace the dike from the foothills up the range front to the crest of the Sierra Nevada (Fig. 26). Here, in the foothills, the dike is shattered and dismembered due to range-front faulting and landsliding. However, most of the large boulders mantling the slope are distinctive porphyritic Golden Bear dike. West of the Independence fault (Moore, 1963), the dike is intact, strikes nearly E-W, dips steeply, and is 10–15 m wide. Here we will see large float boulders of the dike and gain an appreciation for its width the last time it is exposed passing east into Owens Valley. The Golden Bear dike is a K-feldspar quartz monzonite porphyry with euhedral zoned phenocrysts of K-feldspar that range 2–4 cm in length and commonly display Carlsbad twins.
Figure 26. Golden Bear dike on the south side of Pinyon Creek drainage, eastern flank of Sierra Nevada, viewed looking west from Onion Valley Road. Dashed lines follow the outcrop trace of the dike. The prominent peak, Mount Bradley (13,289 ft, 4050 m), is carved in the 102 Ma Bullfrog leucogranite, another tie point across Owens Valley. Stop 8 is along the range front just left (south) of photo.
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Plagioclase is subhedral, subequant, and typically 1–4 mm in diameter. Quartz typically occurs as distinctive equant, euhedral, and bipyramidal crystals 2–5 mm in diameter. Mafic minerals include both biotite and rare hornblende. Stop 9: Red Mountain Fault at Birch Creek Directions. This stop is located at UTM 384884E, 4103959N. From Big Pine, head south on US-395 ~5 mi to North Fish Springs Road. Turn west (right) on North Fish Springs Road and continue until it begins to bend back to the east (toward US395). At the bend (a large triangular intersection), head west on Tinemaha Road. Bear to the right after 300 m, at the base of the Poverty Hills. Birch Creek Road will appear as a dirt road on the right on a very gentle right-hand turn (almost straight) parallel to the creek. This road passes through a cluster of homes, crosses Birch Creek, and then turns back to the west to parallel the creek again. Continue up the switchbacks onto the alluvial fan surface. Stay to the left (straight) at the intersection with a power-line road. The road then veers to the northwest away from Birch Creek toward the granitic Fish Springs Hills. Take the first left, following Birch Creek Road back to the southwest (toward Birch Creek). In ~0.6 mi, the road descends at a W-facing scarp. Park just west of the scarp at UTM 384964E, 4103425N. Walk ~400 m north along the scarp to the large gully. Description. Although recent extension across the southern Owens Valley, near the town of Lone Pine, is primarily accomplished by slip on the Sierra Nevada frontal fault (Le et al., 2007), the geometry and distribution of recent faulting changes markedly in the northern part of the valley (Fig. 1). North-striking normal faults occur in distributed arrays that extend from the foot of the range across the western piedmont to the Owens Valley fault. The best known of these structures is the Fish Springs fault, a W-side-down normal fault that forms the large scarp ~2 km east
of this location. Slip on this structure appears to have been downdip and relatively steady at throw rates of 0.2–0.3 mm/yr (Martel et al., 1987; Zehfuss et al., 2001) since ca. 330 ka. The association of this structure with the surface trace of the Owens Valley fault has led most workers to consider them part of the same fault zone. However, the broad array of normal faults of similar orientation north and west of this location suggests that the Fish Springs fault may be simply one of a distributed array of normal faults that accommodate extension across the Owens Valley at this latitude. The goal of this stop is to examine evidence for the rates of displacement along a couple of these structures. This stop begins at the W-facing scarp of the Red Mountain fault where it displaces alluvial fan deposits of Birch Creek (presently deeply incised to the south of the parking spot; Fig. 27). The topographic expression of the fault is not particularly impressive (a few meters), due to the fact that fan aggradation during and shortly after the last glacial maximum (Zehfuss et al., 2001) on the downthrown hanging-wall block has largely filled the accommodation space. Just north of the road, bouldery surfaces of this fan spill over the scarp and are continuous with sites to the east, dated by Zehfuss et al. (2001) at 13–15 ka (Fig. 27). Approximately 100 m north of the Birch Creek road, two small (~1 m) scarps are present within this alluvial surface. These likely represent the most recent rupture along the Red Mountain fault, and, if representative of average slip, would imply single event throw rates of ~0.2 mm/yr. The Red Mountain fault also displaces lava flows along the southwestern flank of Crater Mountain. Continue northward along the fault scarp, until you reach a large gully draining through the Fish Springs Hills (~300 m). East of this gully, flows are exposed capping a fan deposit with a moderately well-developed soil profile. West of the gully, the flow has been displaced vertically ~12–14 m and buried by younger alluvium. Cosmogenic 36Cl
Figure 27. Field photograph taken looking toward the northeast at a scarp of the Owens Valley fault zone at Birch Creek in Owens Valley (Stop 9). Ridges in the background are the Fish Springs Hills.
Active tectonics of the eastern California shear zone ages from the flow surface indicate that the flow is 70 ± 14 ka (Kirby et al., 2008) and suggest late Pleistocene throw rates of ~0.2 mm/yr. Thus, the Red Mountain fault appears to exhibit similar slip rates to the Fish Springs fault. From this vantage, one can see the Birch Mountain fault, the northernmost segment of the Sierra Nevada frontal fault system. The scarp is apparent south of Tinemaha Creek, where it displaces steep talus and debris-cones at the base of the range, and north of the creek, as it trends upslope toward the prominent cliff at the base of Birch Mountain. The fault displaces Tiogaage moraines in both the Tinemaha and Birch Creek drainages (Fig. 17). Total throw on the structure in these localities varies from 7 to 9 m, and 36Cl ages of boulders from the moraine crest in Tinemaha Creek indicate an age of 13–15 ka (Greene et al., 2007). Thus, throw rates on this segment of the Sierra Nevada frontal fault are ~0.5–0.7 mm/yr Notably, these rates are two to three times greater than those measured along the southern segments of this fault system (Le et al., 2007). Day 3: Big Pine, Furnace Creek, Indian Creek
Stop 10: Owens Valley Fault at Big Pine Dump Directions. This stop is located at UTM 385102E, 4111923N. Just before entering Big Pine from the south on US-395, turn west (left) on Big Pine Dump Road. At the trash/recycling facility, stay to the left onto the dirt road. Stay straight (west) until the intersection with a power-line road. From here, you will see the old town dump. Drive just past the old dump and park under the power lines (~0.3 mi from start of dirt road). Watch out for old nails. Description. At this stop, we will examine evidence for long-term lateral displacement along the Owens Valley fault. From the parking lot near the recycling center, walk west along
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the dirt road for ~300 m. The high-tension power lines visible to the west mark the approximate trace of the Owens Valley fault through this locality. The E-facing scarp of the fault within basalt flows from Crater Mountain is visible as a ~5–7 m high, broken cliff (Fig. 28). At the base of this scarp, fine-grained alluvium has filled a small extensional step in the fault zone. Anthropogenic activity (the historic town dump) has obliterated any trace of the 1872 rupture through this alluvium, although a strand of the fault is visible as a small scarp ~300 m north of this locality (Beanland and Clark, 1994). The apparent right-lateral separation of the flow margin with the modern alluvium is interpreted to reflect dextral displacement across the fault (Kirby et al., 2008). Subsurface surveys using ground-penetrating radar do not reveal any shallow reflectors associated with a buried flow margin on the east side of the fault. Rather, the flow margin visible on the surface appears to represent the former extent of the flow (Kirby et al., 2008). West of the fault, however, young alluvial material did bury the flow margin; prominent islands of basalt are visible above alluvium (Fig. 16). Surveys in this region suggest that the flow margin extends to a position near the prominent protrusion of basalt along the fault scarp (Fig. 16). Restoration of this margin with the exposed flow on the east side of the fault suggests ~235 ± 15 m of lateral displacement along the Owens Valley fault in the past 70 ± 14 k.y. Thus, right-lateral slip rates along the northern Owens Valley fault appear to have been 3.6 ± 1.0 mm/yr over the past 56–80 k.y. (Kirby et al., 2008). Stop 11: Fish Lake Valley Fault at Furnace Creek Directions. This stop is located at UTM 410935E, 4158364N. From Big Pine, take CA-168 east over Westgaard Pass. Continue on CA-168 through Deep Springs Valley to Fish Lake Valley. At the intersection of CA-168 and CA-266 in Oasis, turn north on
Figure 28. Photograph taken looking to the west at fault scarps of the Owens Valley fault zone cutting basalt on the northeast side of Crater Mountain near the Big Pine town dump (Stop 10). Peaks along the eastern Sierra Nevada are visible beyond the scarps.
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Figure 29. View to the northwest along the Death Valley–Fish Lake Valley fault zone at Furnace Creek in central Fish Lake Valley (Stop 11). Note the two prominent fault scarps cutting the Furnace Creek alluvial fan. White Mountain Peak is the prominent summit along the skyline ridge.
CA-266 and go ~7.1 mi to a dirt track leading off to the west (just north of White Wolf Canyon Road to the east). Go ~1.5 mi toward the large shutter-ridge. Continue on the dirt track, bearing to the northwest and paralleling the fault, for another 0.3 mi and park near the incised channel (Fig. 18). Walk ~0.1 mi west up the westernmost fault scarp to the apex of the fan. Description. Just south of the canyon mouth at Furnace Creek, the Fish Lake Valley fault is exposed as two parallel NWstriking strands displacing a late Pleistocene alluvial fan complex (Fig. 29; Frankel et al., 2007b; Reheis et al., 1995; Reheis and Sawyer, 1997). Prominent scarps are exposed in the alluvial fans with the western strand dipping to the NE and the eastern strand dipping to the SW, and with other small normal fault scarps scattered across the fan surfaces (Reheis et al., 1995). The area between the two fault strands is down-dropped in a small pullapart basin (parking area is at the northern end of this pull-apart structure). Just south of the offset late Pleistocene fan is a large NW-SE–striking shutter-ridge, which is bounded on its northeast and southwest sides by the two fault strands (Reheis et al., 1995). Although a number of alluvial fan surfaces are mapped in this location (Reheis et al., 1995), the primary surface of interest is the late Pleistocene Qfio deposit (Qfi of Reheis et al., 1995), which is displaced by the fault (Fig. 18). Cosmogenic nuclide 10 Be dates from the Qfio surface yield an age of 94 ± 11 ka (Frankel et al., 2007b). A recent reexamination of offset channels at this location using ALSM data revised the late Pleistocene offset to 290 ± 20 m (Fig. 18; Frankel et al., 2007b). Combining the 290 ± 20 m of displacement determined from ALSM topography with the cosmogenic 10Be age of 94 ± 11 ka for the Qfio surface yields a minimum late Pleistocene, right-lateral slip rate of 3.1 ± 0.4 mm/yr for the Fish Lake Valley fault zone at Furnace Creek (Frankel et al., 2007b).
Stop 12: Fish Lake Valley Fault at Indian Creek Directions. This stop is located at UTM 396137E, 4182771N. From Dyer, Nevada, head ~8.5 mi north on NV-264 (~16.8 mi north of the California-Nevada border). Turn west on Indian Creek Road (dirt) and continue for ~5 mi until you reach the prominent E-facing scarp near the canyon mouth at UTM 395992E, 4183225N. Park just west of the scarp on the south side of the road (Fig. 19). Walk ~0.3 mi south along the base of the scarp to the third deeply incised channel cutting the footwall. Climb up to the top of the scarp at this location for a good view of the fault zone. Description. The fault scarps at Indian Creek are located at the northern end of the northern Death Valley–Fish Lake Valley fault system. The fault zone splays into numerous normal faults in this location (Reheis et al., 1993), however, a significant strikeslip component is still present (Fig. 19; Frankel et al., 2007b; Reheis et al., 1993; Reheis and Sawyer, 1997). Most of the faulting at this site displaces the late Pleistocene Qfiy surface (Figs. 19 and 30; Qfi of Reheis et al., 1993). The dextral component of slip at this location is restricted to a single strand of the fault near the eastern range-front of the White Mountains (Fig. 19). The strikeslip component of the fault zone is expressed as a prominent scarp cutting the Qfiy and Qfl surfaces (Frankel et al., 2007b; Reheis et al. 1993; Reheis and Sawyer, 1997). The Qfiy surface at Indian Creek has a similar set of soil and morphologic characteristics as the Qfio surface at Furnace Creek. Eight tightly clustered cosmogenic nuclide 10Be surface exposure dates from boulders on the Qfiy fan surface have a mean age and standard deviation of 71 ± 8 ka (Frankel et al., 2007b). Frankel et al. (2007b) used ALSM data to revise the late Pleistocene displacement history to 178 ± 20 m on the basis of six offset channels (Fig. 19). A late Pleistocene slip rate of 2.5 ± 0.4 mm/yr results from the offset determined with ALSM
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Figure 30. Photograph looking southwest along the right-lateral oblique fault scarp at Indian Creek in Fish Lake Valley at the northern end of the Death Valley–Fish Lake Valley fault zone (Stop 12). The eastern White Mountains range front is visible beyond the scarp. Numerous normal faults displace the distal portion of the Indian Creek; all dextral deformation at this location is accommodated on the pictured fault.
data combined with the 71 ± 8 ka 10Be age of the Qfiy surface at Indian Creek (Frankel et al., 2007b). ACKNOWLEDGMENTS Research presented in this guidebook was supported by the National Science Foundation (NSF), the Geothermal Program Office of the China Lake Naval Air Warfare Center, the Southern California Earthquake Center (SCEC), the National Aeronautics and Space Administration, Lawrence Livermore National Laboratory, The Geological Society of America, and the University of California White Mountain Research Station. SCEC is funded by NSF Cooperative Agreement EAR-0106924 and U.S. Geological Survey Cooperative Agreement 02HQAG0008. This is SCEC contribution 1131. ALSM data were collected by the National Center for Airborne Laser Mapping at the University of Florida. A number of people contributed to the work presented in this field guide, including S. Briggs, D. Burbank, N. Dawers, J. Helms, E. Hauksson, J. Hoeft, B. Miller, M. Reheis, M. Rogers, G. Roquemore, T. Sheehan, D. Slemmons, and R. Whitmarsh. Thoughtful reviews by Ernie Duebendorfer and Nathan Niemi helped improve the clarity and presentation of this field trip guidebook. REFERENCES CITED Argus, D.F., and Gordon, R.G., 1991, Current Sierra Nevada–North America motion from very long baseline interferometry: implications for the kinematics of the western United States: Geology, v. 19, p. 1085–1088, doi: 10.1130/0091-7613(1991)0192.3.CO;2. Argus, D.F., and Gordon, R.G., 2001, Present tectonic motion across the Coast Ranges and San Andreas fault system in central California: Geological Society of America Bulletin, v. 113, p. 1580–1592, doi: 10.1130/00167606(2001)1132.0.CO;2. Atwater, T., 1989, Plate tectonic history of the northeast Pacific and western North America, in Winterer, E.L., Hussong, D.M., and Decker, R.W., eds.,
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Stevens, C.H., and Greene, D.C., 1999, Stratigraphy, depositional history, and tectonic evolution of Paleozoic continental-margin rocks in roof pendants of the eastern Sierra Nevada, California: Geological Society of America Bulletin, v. 111, p. 919–933, doi: 10.1130/0016-7606(1999)111 2.3.CO;2. Stevens, C.H., Stone, P., Dunne, G.C., Greene, D.C., Walker, J.D., and Swanson, B.J., 1997, Paleozoic and Mesozoic evolution of east-central California: International Geology Review, v. 39, p. 788-829. Stevens, C.H., Stone, P., Dunne, G.C., Greene, D.C., Walker, J.D., and Swanson, B.J., 1998, Paleozoic and Mesozoic evolution of east-central California, in Ernst, W.G., and Nelson, C.A., eds., Integrated Earth and environmental evolution of the southwestern United States, The Clarence A. Hall, Jr., Volume: Boulder, Colorado, Geological Society of America International Book Series 1, p. 119–160. Stevens, C.H., Stone, P., and Greene, D.C., 2003, Correlation of Permian and Triassic deformations in the western Great Basin and eastern Sierra Nevada: Evidence from the northern Inyo Mountains near Tinemaha Reservoir, east-central California: Geological Society of America Bulletin, v. 115, p. 1309–1311, doi: 10.1130/B25385D.1. Stewart, J.H., 1967, Possible large right-lateral displacement along fault and shear zones in the Death Valley–Las Vegas area, California and Nevada: Geological Society of America Bulletin, v. 78, p. 131–142, doi: 10.1130/0016-7606(1967)78[131:PLRDAF]2.0.CO;2. Strane, M.D., 2007, Slip rate and structure of the nascent Lenwood fault zone, Eastern California [M.S. Thesis]: Chapel Hill, University of North Carolina, 55 p. Stockli, D.F., Dumitru, T.A., McWilliams, M.O., and Farley, K.A., 2003, Cenozoic tectonic evolution of the White Mountains, California and Nevada: Geological Society of America Bulletin, v. 115, p. 788–816, doi: 10.1130/0016-7606(2003)1152.0.CO;2. Sullivan, W.A., and Law, R.D., 2007, Deformation path partitioning within the transpressional White Mountain shear zone, California and Nevada: Journal of Structural Geology, v. 29, p. 583–599, doi: 10.1016/j.jsg.2006.11.001. Thatcher, W., Foulger, G.R., Julian, B.R., Svarc, J., Quilty, E., and Bawden, G.W., 1999, Present-day deformation across the Basin and Range province, western United States: Science, v. 283, p. 1714–1718, doi: 10.1126/ science.283.5408.1714. Thompson, R.A., Milling, M.E., Fleck, R.J., Wright, L.A., and Roger, N.W., 1993, Temporal, spatial, and compositional constraints on volcanism associated with large-scale crustal extension in central Death Valley, California: Eos (Transactions, American Geophysical Union), v. 74, p. 624. Troxel, B.W., 1994, Right-lateral offset of ca. 28 km along a strand of the southern Death Valley fault zone, California: Geological Society of America Abstracts with Programs, v. 26, no. 6, p. 99. Unruh, J.R., Hauksson, E., Monastero, F.C., Twiss, R.J., and Lewis, J.C., 2002, Seismotectonics of the Coso Range–Indian Wells Valley region, California: Transtensional deformation along the southeastern margin of the Sierran microplate, in Glazner, A.F., Walker J.D., and Bartley, J.M., eds., Geologic evolution of the Mojave Desert and Southwestern Basin and Range: Geological Society of America Memoir 195, p. 277–294. Unruh, J.R., Humphrey, J., and Barron, A., 2003, Transtensional model for the Sierra Nevada frontal fault system, eastern California: Geology, v. 31, p. 327– 330, doi: 10.1130/0091-7613(2003)0312.0.CO;2. Unruh, J.R., Monastero, F.C., and Pullammanappallil, S.K., 2008, The nascent Coso metamorphic core complex, east-central California: Brittle upper plate structure revealed by reflection seismic data: International Geology Review (in press). Vines, J.A., 1999, Emplacement of the Santa Rita Flat pluton and kinematic analysis of cross-cutting shear zones, eastern California [M.S. Thesis]: Blacksburg, Virginia Polytechnic Institute and State University, 89 p. Walker, J.D., and Whitmarsh, R.W., 1998, A tectonic model for the Coso geothermal area: U.S. Department of Energy Proceedings Geothermal Program Review XVI, April 1–2, Berkeley, California, p. 2-17–2-24. Weldon, R., Scharer, K., Fumal, T., and Biasi, G., 2004, Wrightwood and the earthquake cycle: What a long recurrence record tells us about how faults work: GSA Today, v. 14, no. 9, p. 4–10, doi: 10.1130/10525173(2004)0142.0.CO;2. Wells, S.G., McFadden, L.D., and Dohrenwend, J.C., 1987, Influence of late Quaternary climatic changes on a desert piedmont, eastern Mojave desert, California: Quaternary Research, v. 27, p. 130–146, doi: 10.1016/00335894(87)90072-X. Wernicke, B., Axen, G.J., and Snow, J.K., 1988, Basin and Range extensional tectonics at the latitude of Las Vegas, Nevada: Geological Society of America
Active tectonics of the eastern California shear zone Bulletin, v. 100, p. 1738–1757, doi: 10.1130/0016-7606(1988)1002.3.CO;2. Wernicke, B., Davis, J.L., Bennett, R.A., Normandeau, J.E., Friedrich, A.M., and Niemi, N.A., 2004, Tectonic implications of a dense continuous GPS velocity field at Yucca Mountain, Nevada: Tectonics, v. 109, doi: 10.1029/2003JB002832. Wesnousky, S.G., 2005, The San Andreas and Walker Lane fault systems, western North America: transpression, transtension, cumulative slip and the structural evolution of a major transform plate boundary: Journal of Structural Geology, v. 27, p. 1505–1512, doi: 10.1016/j.jsg.2005.01.015. Wesnousky, S.G., and Jones, C.H., 1994, Oblique slip, slip partitioning, spatial and temporal changes in the regional stress field, and the relative strength of active faults in the Basin and Range, western United States: Geology, v. 22, p. 1031–1034, doi: 10.1130/0091-7613(1994)0222.3.CO;2. Whistler, D.P., and Burbank, D.W., 1992, Miocene biostratigraphy and biochronology of the Dove Spring Formation, Mojave Desert, California, and characterization of the Clarendonian mammal age (late Miocene) in California: Geological Society of America Bulletin, v. 104, p. 644–658, doi: 10.1130/0016-7606(1992)1042.3.CO;2.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 18 JANUARY 2008
Printed in the USA
The Geological Society of America Field Guide 11 2008
Ediacaran and early Cambrian reefs of Esmeralda County, Nevada: Non-congruent communities within congruent ecosystems across the Neoproterozoic–Paleozoic boundary Stephen M. Rowland Department of Geoscience, University of Nevada, Las Vegas, Nevada 89154-4010, USA Lynn K. Oliver Inyo National Forest, 798 North Main Street, Bishop, California 93615, USA Melissa Hicks ExxonMobil Upstream Research Company, P.O. Box 2189, GW3 966B Houston, Texas 77252-2189, USA
ABSTRACT Esmeralda County, Nevada, is extraordinary for the presence of Ediacaran and early Cambrian reefs at several stratigraphic positions. In this road log and field guide we present descriptions and interpretations of the most instructive exposures of three of these reef-rich intervals: (1) the Mount Dunfee section of the Middle Member of the Deep Spring Formation (Ediacaran in age), (2) the Stewart’s Mill exposure of the Lower Member of the Poleta Formation (mid-early Cambrian), and (3) an exposure on the north flank of Slate Ridge of reefs near the top of the Harkless Formation (latest early Cambrian). We introduce the term “congruent ecosystems” for ecosystems of different age that occupied similar environments. The Ediacaran reefs of the Deep Spring Formation and the early Cambrian reefs of the Lower Member of the Poleta Formation occupied similar environments but exhibit distinctively different ecological structure. Thus we propose these two reef complexes as our premier example of non-congruent communities within congruent ecosystems. Keywords: Cambrian reefs, Ediacaran reefs, congruent ecosystems, archaeocyaths, stromatolites INTRODUCTION
was such a favorable locality for reef development at that time doubtless involves a combination of paleogeographic factors, including latitude, ocean currents, wind currents, and the configuration of the shelf margin (Fig. 1). The following four stratigraphic units all contain reef-rich intervals: (1) the Middle Member of the Deep Spring Formation, (2) the Montenegro Member of the Campito Formation,
Esmeralda County, Nevada, USA, is an extraordinary region for examining reefs of Ediacaran and early Cambrian age. Probably nowhere else in the world has a better representation of accessible, well-preserved, well-exposed reefs that formed at multiple stratigraphic positions within this time interval. The reason this
Rowland, S.M., Oliver, L.K., and Hicks, M., 2008, Ediacaran and early Cambrian reefs of Esmeralda County, Nevada: Non-congruent communities within congruent ecosystems across the Neoproterozoic–Paleozoic boundary, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 83–100, doi: 10.1130/2008.fld011(04). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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Ediacaran and Early Cambrian reefs of Esmeralda County
Figure 1. Paleogeographic setting of the Ediacaran and early Cambrian reefs of Esmeralda County. Modified from Hicks (2006a).
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(3) the Lower Member of the Poleta Formation, and (4) the Harkless Formation (Fig. 2). Of these four “reefy” intervals, the best-known and most thoroughly studied examples of three of them occur quite close to one another in Esmeralda County. The exception is the Montenegro Member of the Campito Formation; the best-known Montenegro reef occurs in the White Mountains of Inyo County, California (Fuller, 1976; Morgan, 1976; Zhou, 1995), which makes it impractical for us to examine on this short field trip. During the past half century, the Ediacaran and Cambrian strata of Esmeralda County, Nevada, and adjacent Inyo County, California, have played an important role in the sorting out of biological and sedimentological events across the ProterozoicPhanerozoic boundary, within North America and globally. This history was recently summarized by Rowland and Corsetti (2002). A field trip to two of the stops visited on this trip (Stops 1 and 2.2), as well as to stops not included in this field guide, was conducted in 2005 by Anderson et al. (2005), with an emphasis on microbialites. Their guidebook is recommended for researchers interested in the reefs described in this field guide, and in microbialites of the Neoproterozoic and Cambrian in general. OBJECTIVES OF THIS FIELD TRIP The objectives of this field trip are (1) to provide an opportunity for participants to examine reefs of the Deep Spring, Poleta, and Harkless Formations; (2) to stimulate discussion among participants—at the outcrop and around the campfire— about reasons for similarities and differences in these reefs; and (3) to explore with participants the concept of “congruent ecosystems” on opposite sides of the Proterozoic-Phanerozoic boundary. We also invite discussion about the question of why the early Cambrian experiment in reef-building by metazoans
was so short lived. The consortium of reef-building archaeocyaths and calcimicrobes existed on Earth for just 11 m.y., from the base of the Tommotian Stage to the top of the Toyonian Stage (Fig. 3). This was followed by an interval of ~40 m.y. (middle Cambrian through the Early Ordovician), during which virtually no metazoan-built reefs developed anywhere on Earth. Several hypotheses have been proposed for this metazoan-reef– free interval (summarized in Rowland and Shapiro, 2002). To us, the most attractive hypotheses are those that involve phenomena associated with global climate change, but these are difficult to rigorously test. We are hoping for a lively discussion of this topic on this trip. WHAT DO WE MEAN BY “CONGRUENT ECOSYSTEMS”? In this field guide we use the term “congruent ecosystems” to characterize ecosystems that occupy the same suite of environments at different times. This is an extension of the term “congruent communities,” originally used by Walker and Laporte (1970). Below is a brief discussion of the history of this terminology and our use of it. Walker and Laporte (1970) coined the term “congruent fossil communities” for fossil communities of different age that occupied similar environments and contain taxa which, although not necessarily closely related taxonomically, have similar autecological characteristics. Their specific examples were four communities that occupied supratidal, high intertidal, low intertidal, and subtidal carbonate environments in the Ordovician Black River Group of New York and four similar communities that occupied the same environments in the Devonian Manlius Formation, also of New York. Sheehan (1996) adapted the concept of congruent communities to Boucot’s (1983) concept of Ecologic
Ediacaran and Early Cambrian reefs of Esmeralda County
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Wyman 0m Figure 2. Stratigraphic column of Ediacaran and Lower Cambrian of the White-Inyo Range and Esmeralda County region. Modified from Hicks (2001).
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Figure 3. Spindle diagrams showing diversity of reef-building taxa in the Ediacaran and Cambrian. From Rowland and Hicks (2004).
Evolutionary Units (EEUs). He also renamed Boucot’s EEUs and defined the fauna within each EEU as an “Evolutionary Fauna,” which he characterized as having had “community stasis” during the EEU in which it occurred. As exemplified by the Ordovician and Devonian examples of Walker and Laporte (1970), a community within any of Sheehan’s Evolutionary Faunas is inferred to have had a similar ecologic structure to a community in a similar environmental setting in a different EEU. In Sheehan’s (1996) terminology, the Cambrian period has two EEUs, which he called C1 and C2. C1 corresponds to the portion of the early Cambrian in which archaeocyaths occur (Tommotian through Toyonian Siberian Platform Stages; Fig. 3). On this field trip, this EEU includes the Poleta and Harkless Formations. Sheehan (1996) suggested that it may be necessary to erect another EEU in the pre-trilobite portion of the Cambrian, the Nemakit-Daldynian interval of the Siberian Platform (Fig. 3). He also proposed an EEU for the late Neoproterozoic, when the Ediacaran fauna lived. He named this unit E1. On this field trip, we will examine reefs that occur in Sheehan’s EEUs E1 and C1. However, the structure of the reefbuilding communities in these two intervals is very different. In contrast to the similarities observed between EEUs within the Phanerozoic eon, such as the Ordovician and Devonian examples of Walker and Laporte (1970), a major point of this field trip will be that reef communities across the EdiacaranCambrian boundary exhibit very different ecological structure, although they lived in comparable environments. For this reason, we refer to these as “non-congruent communities” within “congruent ecosystems.”
STOP 1: THE STROMATOLITE REEF COMPLEX OF THE MIDDLE DEEP SPRING AT MOUNT DUNFEE (EDIACARAN IN AGE) At Mount Dunfee, near Gold Point, Nevada, the Middle Member of the Deep Spring Formation is unusually rich in stromatolites and other microbialites. The Neoproterozoic-Cambrian boundary at this locality, as defined by the lowest occurrence of Treptichnus pedum, occurs high within the Middle Member, above the highest of the microbialite horizons. This section thus provides an excellent opportunity to examine morphological diversity in latest Neoproterozoic reefs. Figure 4 is a stratigraphic column of this section, while Figure 5 shows two photographs of the section. We recognize three depositional systems within the Middle Member Deep Spring Formation at Mount Dunfee. The first is a siliciclastic intertidal and shallow subtidal system that is represented by quartz siltstone and sandstone in the lower 42 m of the 175m-thick section. These sediments were deposited in tide- and storm-dominated subtidal and intertidal environments prior to the initiation of carbonate sediment on the miogeoclinal ramp. Analysis of herringbone cross-stratification within the dolo-allochemic quartz sandstone shows a bimodal pattern that records tidal flow perpendicular to the paleo-shelf margin. Overlying these siliciclastic-dominated sediments are oolites and microbialites, representing peritidal reefs and shoals. Strata of this carbonate-dominated depositional system occupy two intervals within the section. They cap the siliciclastic-dominated interval at the bottom of the section, and they also occupy most of the
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Figure 4. Stratigraphic column of the Middle Member of the Deep Spring Formation at Mount Dunfee. From Oliver (1990).
upper half of the section. In each case, these reef-shoal sediments complete a shoaling-upward cycle within the section. The third depositional system is a mixed siliciclastic-carbonate shallow subtidal system, represented by a 40-m-thick interval of micritic sandstone roughly in the middle of the section. These sediments record the early phase of a second shoaling-upward cycle within the Middle Member, during a time of greater carbonate production. Figure 6 is a stratigraphic column of a portion of the Middle Member, showing microbial horizons A through E and adjacent lithologies. The microbialites are morphologically and structurally diverse, including bioherms of digitate stromatolites, bioherms and biostromes of inclined columnar stromatolites, isolated massive and hemispheroidal stromatolites, stromatolitic thrombolites, and a biostrome of cryptomicrobial boundstone. Some of these were microbial reefs that had topographic relief, formed in active agitated waters, and exerted a physical control over their environment. A few stromatolitic thrombolites occur in this section. This locality is the oldest known occurrence of thrombolites in southwestern North America, although considerably older
examples are known from northwestern Canada. Figure 7 shows examples of stromatolites, thrombolites, and oolites. Meter-scale cyclicity is conspicuous within the portion of the section that contains microbial horizons C, D, and E (Fig. 8). The Mount Dunfee microbialites incorporate up to 32 wt% detrital quartz within their microstructures. They represent an intermediate form between “pure” siliciclastic and “pure” carbonate stromatolites—theoretical end members in a continuum. An important prerequisite for preserving microbialite fabrics in the stratigraphic record is early carbonate cementation. The well-preserved, quartz-rich Mount Dunfee microbialites provide a datum concerning the amount of siliciclastic material a microbialite can contain without a substantial loss of preservation potential and morphological variability. Possible close modern analogs of some of the Mount Dunfee stromatolites occur in subtidal, current-swept channels in the Bahamas. Like their ancient counterparts in the Deep Spring Formation, the Bahamian stromatolites are predominantly columnar, they are inclined into the flood tide, and they are closely associated with oolite sand waves that episodically bury them.
Figure 5. (A) View toward the east of the Middle Member Deep Spring section at Mount Dunfee; ~150 m of strata are in view, and beds dip ~40°NE. The Ediacaran-Cambrian boundary is located a few tens of meters below the top of the Middle Member, above Microbial Horizon I, based on the lowest occurrence of the trace fossil Treptichnus pedum at that horizon. (B) View toward the southeast of a portion of the Middle Deep Spring containing microbial horizons A through E. Compare with Figure 6. Strata within the white box are ~12 m thick. Modified from Oliver and Rowland (2002).
Figure 6. Detailed stratigraphic column of the interval shown in Figure 5B. From Oliver (1990).
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Figure 7. Photographs of exposures in the Mount Dunfee section. (A) Vertical cliff-face exposure of inclined columnar stromatolites; lower microbial interval. (B) Vertical cliff-face exposure of inclined columnar stromatolites; lower microbial interval; irregular shapes of stromatolites is caused by close packing; width of photo is ~35 cm. (C) Inclined columnar stromatolite with parallel margins; lower microbial interval; pocket knife in center of photo is 9 cm long. (D) Lynn Oliver sitting on Bioherm C.2, a lenticular bioherm of inclined stromatolites. (E) Bioherm C.2, a lenticular bioherm of inclined stromatolites; hammer for scale. (F) Columnar stromatolites below, abruptly overlain by oolite; Microbial Horizon D. (G) Stromatolitic thrombolite, which has a mixture of laminated and clotted fabrics; Microbial Horizon H. From Oliver and Rowland (2002).
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Figure 8. Field sketch of five meter-scale cycles within microbial horizons C, D, and E. Each cycle begins with a stratiform stromatolite, which usually grades abruptly into columnar stromatolites. In cycle 3, oolite grainstone interrupts the normal succession. We interpret these cycles to be deepening upward cycles.
As represented schematically in Figure 9, we suggest that the Mount Dunfee section, with its conspicuous and episodically persistent peritidal reef and shoal facies, represents the nucleation of carbonate sedimentation in each of two shoaling-upward, carbonate-capped cycles within the Middle Member Deep Spring Formation. From this position on the ramp, carbonate sedimentation expanded seaward and cratonward, ultimately building into a narrow, discontinuous, episodically emergent reef and shoal complex. The Mount Dunfee section thus represents a geographically restricted, temporally persistent locus of carbonate sedimentation and microbialite reef construction during the waning portion of the Ediacaran Period. STOP 2.1: THE EARLY CAMBRIAN ARCHAEOCYATHAN— RENALCIS—THROMBOLITE REEF COMPLEX OF THE LOWER POLETA AT STEWART’S MILL The Stewart’s Mill locality of the Lower Member of the Poleta Formation is the most instructive exposure of the reef facies of this stratigraphic unit. Aspects of this exposure have been described by Rowland (1984), Rowland and Gangloff (1988), and Rowland and Shapiro (2002). The exposure is a pyramidshaped hill (Fig. 10). The lower third is green shale of the Mon-
tenegro Member of the Campito Formation. The middle portion of the hill consists of 70 m of thrombolitic and archaeocyathanRenalcis boundstone of the Lower Member of the Poleta Formation (Fig. 11). Capping the hill is a cliff-forming interval, 56 m thick, of oolite, biostromes, and packstone (Figs. 11 and 12). The interval of main interest for this field trip is the 70 m “reefy” interval below the cliffs. There is a complex mosaic of interfingering facies in this portion of the exposure (Fig. 11), but the basic succession consists of the following facies, in ascending order: (1) bioclastic lime mudstone, (2) thrombolites with sparse archaeocyaths, (3) Renalcis-dominated boundstone, (4) green mudshale with lenses of skeletal wackestone, (5) archaeocyathdominated boundstone, and (6) oolite grainstone. We interpret Facies 4, the green mudshale (which is poorly exposed), to represent a bypass channel by which siliciclastic sediment was shunted across the carbonate shelf margin. Of particular interest within this exposure is the conspicuous ecological zonation that occurs within the succession of reef-lagoon facies. The lower portion of this interval is dominated by thrombolite, with archaeocyaths comprising less than 3% of the volume of the rock (Fig. 13). In stark contrast, the uppermost portion of this interval is dominated by conspicuous branching archaeocyaths, which comprise up to 38% of the volume of the rock (Fig. 14). Figure 15 summarizes the characteristics within each zone. The three most conspicuous changes that occur, from bottom to top within this 65-m-thick interval, are (1) the relative abundance of archaeocyaths increases (Fig. 16); (2) the morphology of archaeocyaths changes, with non-branching forms dominant throughout most of the section and branching forms becoming dominant within the uppermost 30 m (Fig. 14); and (3) decimeter- and meter-scale cavities (preserved as dolomite-rich, orange patches) are absent in the lower one-third of this interval, but are very conspicuous in the upper two-thirds. Such zonation has been described in fossil reefs of a variety of ages, including Ordovician, Silurian, Devonian, and Cretaceous (Walker and Alberstadt, 1975). This locality is the only example of ecological zonation described in a Cambrian reef. Walker and Alberstadt (1975) interpreted such zonation to be a record of ecological succession, but in cases such as this one, where very long time intervals are involved, ecological succession is not the process being recorded in the rocks (see discussion by Rowland and Gangloff, 1988). Rather, we interpret the conspicuous zonation in the Lower Poleta at Stewart’s Mill to represent adjacent facies that migrated over one another during a marine transgression. Using these three trends, Rowland and Shapiro (2002) divided this interval into three environmental zones: (1) a lower back-reef lagoon zone, (2) a middle low-energy reef-crest zone, and (3) a high-energy reef-crest zone (Fig. 15). The presence of conspicuous primary cavities (now filled with partially dolomitized micrite) distinguishes the reef-crest facies from the underlying lagoonal facies, and the presence of abundant, conspicuous branching archaeocyaths distinguishes the high-energy reef-crest zone from the underlying low-energy reef-crest zone.
Ediacaran and Early Cambrian reefs of Esmeralda County
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Portion of Lower Member Wood Canyon Formation
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Base of Middle Member Deep Spring Formation Intertidal & Supratidal Carbonates Ediacaran - Cambrian Boundary (lowest occurrence of Treptichnus pedum)
Figure 9. Schematic regional synthesis of Middle Member Deep Spring facies at Mount Dunfee and correlative strata to the northwest and southeast.
Figure 10. Photograph of the Stewart’s Mill exposure of the Lower Member of the Poleta Formation and uppermost portion of the Montenegro Member of the Campito Formation. Lower third of hillside is shale of the uppermost Campito Formation. Middle third is 70-m-thick interval of predominantly archaeocyathan-Renalcis reef facies, interrupted by a poorly exposed shale and siltstone interval. Cliff-forming upper third of exposure is 60-m-thick interval, predominantly oolite. View is toward the north. Compare with Figure 11. Photo by S. Rowland.
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Figure 11. Lower Poleta facies at the Stewart’s Mill locality; qtz siltst—quartz siltstone. From Rowland (1981).
It is noteworthy that the Lower Poleta reef complex at Stewart’s Mill has abundant microbialites but, in stark contrast to the Deep Spring reefs at Mount Dunfee, there are no stromatolites. All of the microbialite in the Poleta at this locality has a clotted, thrombolitic texture.
Poleta reefs and oolite shoals formed during an interval of carbon isotopic stasis.
Chemostratigraphy of the Lower Poleta
STOP 2.2: PATCH REEFS NEAR THE TOP OF THE HARKLESS FORMATION ON SLATE RIDGE: THE LAST GASP OF METAZOAN REEF-BUILDING IN THE EARLY CAMBRIAN
As part of Hicks’s dissertation research, we sampled this exposure of the Lower Poleta for isotopic analysis, with the hope of using chemostratigraphy to assist in intercontinental correlation. Her results are presented in Figure 17. Although the early Cambrian is known for its carbon isotope excursions, surprisingly this interval showed very little variation in its δ13C record. There is a slight drift in δ13C values ranging from 0‰ to 1‰ at the bottom of the section to −1‰ at the top. The fact that the δ13C values do not covary with the δ18O values (Fig. 17) suggests that the stasis in the δ13C values is a primary signal and not a product of diagenetic alteration (Hicks, 2006a). Apparently, the Lower
The final stop on this field trip will be in the upper portion of the Harkless Formation to examine some meter-scale patch reefs studied by Hicks (2001) (Fig. 18). These Harkless reefs are among the youngest Cambrian reefs in the world in which metazoans played a significant constructional role. They are approximately equivalent in age to the well-studied reefs of the Forteau Formation of eastern Canada. Following the disappearance of these late early Cambrian archaeocyath-rich reefs, microbialites (stromatolites, thrombolites, and dendrolites) became the prevailing reefs through the middle and late Cambrian (Furongian) and Early Ordovician (Rowland and Shapiro, 2002).
Figure 12. Stratigraphic column of the oolite shoal facies complex at the Stewart’s Mill locality. Zone numbers refer to inferred amount of agitation in the depositional environment; zone 1 is constant agitation, while zone 4, at the other extreme, represents only occasional, storm-generated agitation. From Rowland (1981).
Figure 14. Abundant branched archaeocyaths high in the Lower Poleta boundstone interval at Stewart’s Mill. Archaeocyaths comprise up to 38% of the volume of the rock at this level. Photo by S. Rowland.
Figure 13. Typical exposure of thrombolite-rich interval low in the Lower Poleta section at Stewart’s Mill. Archaeocyaths are sparse within this facies, comprising less than 3% by volume. This facies corresponds to the back-reef lagoon zone of Figure 15. Rock hammer for scale. Photo by S. Rowland.
Ediacaran and Early Cambrian reefs of Esmeralda County
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Figure 15. Vertical ecological zonation in the reef-lagoon facies complex of the Lower Member of the Poleta Formation at Stewart’s Mill. From Rowland and Hicks (2004).
Figure 16. Facies relationships within the reef-lagoon interval at Stewart’s Mill. Numbers below the black triangles indicate the volumetric abundance of archaeocyaths. From Rowland and Gangloff (1988).
Ediacaran and Early Cambrian reefs of Esmeralda County Stewart's Mill Poleta Formation Section -18
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Figure 17. Values of δ18O and δ13C (‰) for the Lower Poleta at Stewart’s Mill. From Hicks (2006a).
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Reef A Figure 18. Three meter-scale patch reefs near the top of the Harkless Formation at Site 1 (Stop 2.2) of Hicks (2001). View is toward the north. Photo by M. Hicks.
Figure 19. Stratigraphic column of an interval near the top of the Harkless Formation, at Site 1 (Stop 2.2) of Hicks (2001), including a patch reef. From Hicks (2001).
Ediacaran and Early Cambrian reefs of Esmeralda County Figure 19 is a stratigraphic column of the locality we will visit, and Figure 20 shows the sedimentological and paleontological details of the reefs and associated facies. As is the case with the Forteau reefs, these Harkless reefs are associated with a high diversity fauna of dwelling organisms that were important components of community structure but were not involved in reef construction. Figure 21 shows the volumetric percentages of various constituents of the reef framestone. Some reefs within the uppermost Harkless contain one of the earliest reported corals from North America, Harklessia yuenglingensis, which has similarities with tabulate corals (Hicks, 2006b). Unfortunately, the reefs with Harklessia are not easily accessible, so these early Cambrian corals will not be observed on this field trip. ROAD LOG DAY 1: EDIACARAN MICROBIAL REEFS OF THE MIDDLE MEMBER DEEP SPRING FORMATION AT MOUNT DUNFEE This day will involve a drive of ~200 mi (320 km) into Esmeralda County, Nevada, USA. We’ll travel north from Las Vegas on U.S. 95, with a stop for lunch and a restroom break in Beatty, Nevada. The road distance from the University of Nevada–Las Vegas (UNLV) to Beatty is ~125 mi (200 km). The detailed road log begins at the Death Valley Nut and Candy Company in Beatty. We’ll spend about three hours in the afternoon examining the microbial reefs of the Horse Spring Formation. Cumulative mi (km) 0.0
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51.9
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58.9
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66.4 (106.8)
66.5 (106.9) 66.7 (107.3) 67.0 (107.8)
68.2 (109.7) 69.2 (111.3) 69.6 (112.0)
Directions Leave the Death Valley Nut and Candy Company parking lot (and Eddy World gas station) and head north (left) on U.S. 95. Turn left (southwest) on Nevada Highway 266 at Lida Junction. Turn left (southwest) on Nevada Highway 744 toward Gold Point. “Downtown” (Gold St. and 2nd Ave.) Gold Point, a semi-ghost town. Turn left on 2nd Avenue. Bear right on main dirt road (extension of Orleans St.). Turn right at intersection of two dirt roads. Turn left on dirt road. (On the Gold Point 7.5 min Quadrangle map, this is a prominent road that heads southeast around the south side of Mount Dunfee.) Junction; no turn. Junction marked by a rusty old street sign; no turn. Bold ridge of rhyolite on left (the “rhyolite mammoth”); park on left side of road.
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Stop 1: Microbialite Reefs and Associated Sedimentological Features of the Middle Member of the Deep Spring Formation Here we begin a hike of ~1.5 mi, over uneven terrain, into the Mount Dunfee section. The destination is the exposure beneath the second “e” in the word “Dunfee” on the quadrangle map. See detailed description above. Cumulative mi (km)
Directions
69.6 (112.0) Leave the “rhyolite mammoth” and return to Gold Point. 72.8 (117.1) Turn right on Main Street (Mitchell’s Mercantile is a prominent building on Main St.). 72.9 (117.3) Turn left on dirt road that leads north across Lida Valley 80.0 (128.7) Turn right at intersection of two dirt roads. 80.4 (129.4) Intersection with Nevada Highway 266; turn right onto pavement. 81.4 (131.0) Turn left onto well-used dirt road marked by a faded, trapezoid-shaped Bureau of Land Management sign just beyond turn. 84.3 (135.6) Bear left at fork in road. 84.4 (135.8) Turn left onto a track that leads into the hills. 85.0 (136.8) Campsite in abandoned quarry. DAY 2: EARLY CAMBRIAN ARCHAEOCYATHANRENALCIS REEFS OF THE POLETA AND HARKLESS FORMATIONS We’ll visit two localities this morning, both of which are just a few miles from our campsite. The first will be the massive reef and shoal complex at the Stewart’s Mill locality of the Lower Member of the Poleta Formation, and the second will be some meter-scale patch reefs near the top of the Harkless Formation. Cumulative mi (km)
Directions
85.0 (136.8) Leave campsite and retrace route back to Nevada Highway 266. 88.6 (142.6) Turn right (west) onto Nevada Highway 266. 89.6 (144.2) Turn left onto dirt road (marked by a stop sign). 89.9 (144.6) Cross intersection; no turn. 91.0 (146.4) Turn right onto track that leads toward prominent pyramid-shaped hill. 91.1 (146.6) Parking area. Stop 2.1: Stewart’s Mill Locality of the Lower Member of the Poleta Formation It is a short but steep hike up the hill, through the shales of the Montenegro Member of the Campito Formation, and into
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Figure 20. Sedimentological features of the reef-bearing interval of the Harkless Formation at Site 1 (Stop 2.2). From Hicks (2001).
Figure 21. Mean percentages of constituents in Site 1 (Stop 2.2) reefs. “Other” category refers to brachiopods and unidentifiable shell material. Sm. block and lg. block refer to calcite cement. From Hicks (2001).
Ediacaran and Early Cambrian reefs of Esmeralda County the Lower Member of the Poleta Formation. See earlier detailed description of this locality. A Note about Collecting Samples at This Site This exposure of the Lower Poleta is one of the best examples of early Cambrian archaeocyathan reefs in the world. Professors from many universities bring their students here, and paleontologists from all over the world come here to examine this reef complex. If you want to collect samples, please confine your collecting to loose material on the slope, of which there is plenty. Please leave the in situ fossils for future generations of geologists and paleontologists to examine and study in context. Cumulative mi (km) 91.1 91.2 92.3 92.6
(146.6) (146.7) (148.5) (149.0)
101.1 (162.7) 101.7 (163.6)
101.9 (164.0)
103.5 (166.5) 103.8 (167.0)
Directions Retrace route back to well-used dirt road. Turn left on road. Cross intersection; no turn. Junction with Nevada Highway 266. Turn right (east). Turn right (toward Gold Point) onto Nevada Highway 774. As paved road turns to the right, continue straight onto dirt track. (Beginning at this point the route requires high clearance and a low center of gravity; 4-wheel drive recommended.) Cross intersection with another track. No turn. Our route heads toward prominent black peak in the distance. Turn left out of wash onto track that leads up onto desert surface above level of wash. Parking area for Stop 2.2.
Stop 2.2: Patch Reefs of the Harkless Formation Three meter-scale patch reefs of the Harkless Formation are exposed on the low ridge immediately to the north of this parking area. See earlier detailed description of this locality. After examining these reefs, we will eat lunch and then begin our return to Las Vegas. Cumulative mi (km) 103.8 105.9 106.5 113.5 165.4
(167.0) (170.4) (171.4) (182.6) (266.1)
Directions Retrace route back to Nevada Highway 774. Turn right on Nevada Highway 774. Turn right (east) on Nevada Highway 266. Turn right (south) on U.S. 95 toward Beatty. Brief rest stop at the Death Valley Nut and Candy Co. in Beatty. Continue south on U.S. 95 to Las Vegas. Distance from Beatty to UNLV is ~125 mi (200 km). Total distance traveled on this field trip is 415 mi (668 km). End of road log.
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ACKNOWLEDGMENTS We thank Tom Anderson, Ernie Duebendorfer, and Russell Shapiro for helpful reviews of this field guide, and we thank Becki Huntoon, Peter Starkweather, and Ethan Starkweather for help with figures. REFERENCES CITED Anderson, T.B., Hicks, M., and Shapiro, R.S., 2005, Microbialite sediments of the Death Valley area, in Stevens, C., and Cooper, J., eds., Western Great Basin Geology: fieldtrip guidebook and volume prepared for the joint meeting of the Cordilleran Section, GSA, and Pacific Section, AAPG: Fullerton, California, Pacific Section SEPM (Society for Sedimentary Geology), p. 67-107. Boucot, A.J., 1983, Does evolution take place in an ecological vacuum? II: Journal of Paleontology, v. 57, p. 1–30. Fuller, D.R., 1976, Paleoenvironmental analysis of an Early Cambrian archaeocyathid reef in the White-Inyo Mountains, California [M.S. thesis]: Idaho State University, 53 p. Hicks, M., 2001, Paleoecology of upper Harkless archaeocyathan reefs in Esmeralda County, Nevada [M.S. thesis]: Las Vegas, University of Nevada, 150 p. Hicks, M., 2006a, Characterizing global archaeocyathan reef decline in the Early Cambrian: Evidence from Nevada and China [Ph.D. dissertation]: Las Vegas, University of Nevada, 138 p. Hicks, M., 2006b, A new genus of Early Cambrian coral in Esmeralda County, southwestern Nevada: Journal of Paleontology, v. 80, p. 609–615, doi: 10.1666/0022-3360(2006)80[609:ANGOEC]2.0.CO;2. Morgan, N., 1976, The Montenegro bioherms: their paleoecology, relation to other archaeocyathid bioherms and to early Cambrian sedimentation in the White and Inyo Mountains, California, in Moore, J.N., and Fritsche, A.E., eds., Depositional Environments of Lower Paleozoic Rocks in the White-Inyo Mountains, Inyo County, California, Pacific Coast Paleogeographic Field Guide 1: Los Angeles, California, Pacific Section, Society of Economic Paleontologists and Mineralogists, p. 13-17. Oliver, L.K., 1990, Stromatolites of the Middle Member of the Deep Spring Formation, Esmeralda County, Nevada [M.S. thesis]: Las Vegas, University of Nevada, 150 p. Oliver, L.K., and Rowland, S.M., 2002, Microbialite reefs at the close of the Proterozoic Eon: The Middle Member Deep Spring Formation at Mount Dunfee, Nevada, in Corsetti, F.A., ed., Proterozoic-Cambrian of the Great Basin and Beyond: Fullerton, California, Pacific Section SEPM (Society for Sedimentary Geology), p. 97–122. Rowland, S.M., 1981, Archaeocyathid bioherms in the Lower Poleta Formation, Esmeralda County, Nevada, in Taylor, M.E., and Palmer, A.R., eds., Second International Symposium on the Cambrian System, Guidebook for Field Trip 1: Denver, Colorado, p. 44-49. Rowland, S.M., 1984, Were there framework reefs in the Cambrian?: Geology, v. 12, p. 181–183, doi: 10.1130/0091-7613(1984)12 2.0.CO;2. Rowland, S.M., and Corsetti, F.A., 2002, A brief history of research on the Precambrian-Cambrian boundary in the southern Great Basin, in Corsetti, F.A., ed., Proterozoic-Cambrian of the Great Basin and Beyond: Fullerton, California, Pacific Section SEPM (Society for Sedimentary Geology), p. 97–122. Rowland, S.M., and Gangloff, R.A., 1988, Structure and paleoecology of Lower Cambrian reefs: Palaios, v. 3, p. 111–135, doi: 10.2307/3514525. Rowland, S.M., and Hicks, M., 2004, The Early Cambrian experiment in reefbuilding by metazoans, in Lipps, J.H., and Waggoner, B.M., eds., Neoproterozoic-Cambrian biological revolutions: The Paleontological Society Papers, v. 10, p. 107–130. Rowland, S.M., and Shapiro, R.S., 2002, Reef patterns and environmental influences in the Cambrian and earliest Ordovician, in Kiessling, W., Flűgel, E., and Golonka, J., eds., Phanerozoic Reef Patterns: Tulsa, Oklahoma, SEPM (Society for Sedimentary Geology) Special Publication 72, p. 95–128. Sheehan, P.M., 1996, A new look at Ecological Evolutionary Units (EEUs): Palaeogeography, Palaeoclimatology, Palaeoecology, v. 127, p. 21–32, doi: 10.1016/S0031-0182(96)00086-7.
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Walker, K.R., and Alberstadt, L.P., 1975, Ecological zonation as an aspect of structure in fossil communities: Paleobiology, v. 1, p. 238–257. Walker, K.R., and Laporte, L.F., 1970, Congruent fossil communities from Ordovician and Devonian carbonates of New York: Journal of Paleontology, v. 44, p. 928–944.
Zhou, X., 1995, Lower Cambrian bioherms in central Nevada and eastern California [M.S. thesis]: Las Vegas, University of Nevada, 106 p. MANUSCRIPT ACCEPTED BY THE SOCIETY 24 JANUARY 2008
Printed in the USA
The Geological Society of America Field Guide 11 2008
Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert Andrew P. Barth Department of Earth Sciences, Indiana University–Purdue University, Indianapolis, Indiana 46202, USA J. Lawford Anderson Department of Earth Sciences, University of Southern California, Los Angeles, California 90089, USA Carl E. Jacobson Department of Earth and Atmospheric Sciences, Iowa State University, Ames, Iowa 50011, USA Scott R. Paterson Department of Earth Sciences, University of Southern California, Los Angeles, California 90089, USA Joseph L. Wooden U.S. Geological Survey, 345 Middlefield Road, Menlo Park, California 94025, USA
ABSTRACT This field guide describes a two-and-one-half day transect, from east to west across southern California, from the Colorado River to the San Andreas fault. Recent geochronologic results for rocks along the transect indicate the spatial and temporal relationships between subarc and retroarc shortening and Cordilleran arc magmatism. The transect begins in the Jurassic(?) and Cretaceous Maria retroarc fold-andthrust belt, and continues westward and structurally downward into the Triassic to Cretaceous magmatic arc. At the deepest structural levels exposed in the southwestern part of the transect, the lower crust of the Mesozoic arc has been replaced during underthrusting by the Maastrichtian and/or Paleocene Orocopia schist. Keywords: California, structural geology, petrology, geochronology, tectonics OVERVIEW Achieving the goal of understanding the geodynamic evolution of a convergent continental margin arc requires an understanding of the interplay between magmatic and tectonic processes through time. How did shortening in the southwestern North American Cordillera relate in time and space to arc magmatism?
Advances in geochronology applied to regional field and petrologic studies in an exhumed arc and its associated thrust belts in southern California are illuminating the timing of underthrusting and shortening relative to voluminous arc magmatism. Cordilleran foreland shortening and arc magmatism were broadly contemporaneous, but their relative timing remains poorly known. As a result, it has long been argued whether
Barth, A.P., Anderson, J.L., Jacobson, C.E., Paterson, S., and Wooden, J.L., 2008, Magmatism and tectonics in a tilted crustal section through a continental arc, eastern Transverse Ranges and southern Mojave Desert, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 101–117, doi: 10.1130/2008. fld011(05). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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shortening leads to, or is the structural response to precursory magmatic thickening in the arc. As higher precision U-Pb ages become more readily available for more plutons in the Cordilleran arc, it seems likely that batholith emplacement was an episodic phenomenon superimposed on a longer-term background magmatic flux. In a like manner, more precise geochronologic control on evolution of the foreland fold-and-thrust belt is necessary to better establish the orogen-wide stress field over long time scales, and thus clarify whether foreland shortening is a cause or consequence of convergent margin arc magmatism. This field trip offers an overview of the tectonic evolution of a portion of the retroarc region and a tilted section through the crust of the Cordilleran arc. An east to west transect will afford us a view of arc tectonics in the Cretaceous, and a top-down view of variations in the composition and emplacement style of Mesozoic igneous rocks in a tilted arc section. Comparison of the plutonic and shortening records allows us to relate deformation in time and space to the progress of arc magmatism. The tilted arc section was likely created during shallow underthrusting of oceanic lithosphere beneath this long-lived Mesozoic arc, and so the timing of this underthrusting event is key to understanding arc extinction and exhumation. In the Orocopia Mountains, exhumation has exposed detached plates of moderate and deep structural levels of the arc. These are underlain by Orocopia Schist, which is part of the Pelona-Orocopia-Rand schist subduction complex underplated beneath southern California and western Arizona during the Laramide orogeny. Here we will consider both the underplating and exhumation history of the schist, as well as geologic relations within the overlying crystalline and supracrustal rocks of native North America. On this trip, we will follow an east to west transect from the retroarc fold-and-thrust belt along the Arizona–California border, westward into the Mesozoic arc. Along this transect, the timing of Mesozoic plutonism is now reasonably well characterized using U-Pb geochronology, which has also been applied to estimating the timing of shortening in the retroarc and subarc regions. At this latitude, the Mesozoic plutonic record incorporates overlapping segments of Permo-Triassic, Middle Jurassic, Late Jurassic and Late Cretaceous arc segments. As we progress further to the west into the core of the arc, we will see a region of the arc that enjoyed shallow underthrusting and consequent extreme extension during and following the early
Cenozoic Laramide orogeny. The result of this variable extension is regional west-side-up tilting that allows us a view of depth-dependent changes in arc magmatism and processes of magma transport and emplacement. DAY 1 Directions to Stop 1 Depart Blythe, traveling west on Interstate 10 (I-10). Exit Mesa Road [UTM E 0710924 N 3721114 (NAD 27 CONUS)]. Continue west on Black Rock Road to Stop 1 [0708035 3721024]. Stop 1. Mule Mountains Thrust Zone and McCoy Mountains Formation This stop finds us in the south-central part of the McCoy basin, a west-northwest–trending basin in western Arizona and southeastern California. The later structural evolution of the basin is characterized by crystalline thrust sheets of opposing vergence, primarily exposed along the margins of the basin. The basin margins are visible as Middle Jurassic plutonic rocks thrust over Jurassic volcanic rocks in the Mule Mountains visible to the south, and Proterozoic basement and Paleozoic cratonal cover in the Maria Mountains visible to the northeast (Stone, 2006; Fig. 1). At this location, we can see rocks characteristic of the basin and its margins; Jurassic metavolcanic rocks of the Dome Rock sequence are thrust northward over clastic sediments of the Jurassic(?) and Cretaceous McCoy Mountains Formation. The earlier evolution of the McCoy basin is controversial, depending on interpretation of the origin of the 4–7.5 km of clastic sediment of the McCoy Mountains Formation that have yielded very few, problematic fossils. An inferred Jurassic depositional age for the McCoy Mountains Formation led to an early hypothesis that during most of its evolution the basin was bounded by sinistral transtensional faults of the hypothetical Mojave-Sonora megashear (Harding and Coney, 1985; Saleeby and Busby-Spera, 1992; Anderson and Nourse, 2005). Alternatively, an inferred Cretaceous depositional age (Reynolds et al., 1986; Stone et al., 1987; Tosdal, 1990; Tosdal and Stone, 1994) led to the hypothesis that at least some basin sedimentation was associated with shortening in the Maria fold-and-thrust belt.
Figure 1. Simplified cross section (adapted from Stone, 2006) of the McCoy Mountains Formation (gray shaded) in the McCoy basin. Tu—undifferentiated Tertiary; Mg—Mesozoic granitic and gneissic rocks; Jv—Jurassic volcanic rocks; Pz—Paleozoic and Triassic(?) sedimentary rocks; Xg—Proterozoic crystalline rocks.
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Recent petrography, zircon geochronology and geochemistry of McCoy sandstones provide new limits on the timing and origin of McCoy basin fill (Barth et al., 2004). Detrital zircon populations in sandstones vary systematically with stratigraphic height (Fig. 2) and sandstone composition. Minimum detrital zircon ages of 116 and 109 Ma in the lower part of basal sandstone member 2 demonstrate that >90% of the formation was deposited in post-Aptian, middle Early to middle Late Cretaceous time. Comparison of detrital zircon ages to sandstone compositions indicates that sand was derived from both sides of the basin, and that deposition was synchronous with volcanism to the west and shortening and exhumation in the fold-and-thrust belt. These results support the hypothesis that the McCoy basin originated as a retroarc foreland basin during Cretaceous thrusting. This result is significant because the McCoy basin then links shortening and foreland basin sedimentation in the Maria fold-and-thrust belt to contemporaneous retroarc shortening in the southern Sevier belt to the north and the Sonoran thrust belt to the south, recording regional crustal shortening synchronous with voluminous upper crustal arc magmatism. Directions to Stop 2 Depart stop, return east on Black Rock Road to I-10. Continue west on I-10. Exit Eagle Mountain Road [0643527 3730127]. Continue north on Eagle Mountain Road to Stop 2 [0643247 3738425]. Stop 2. Middle Jurassic Eagle Mountain Pluton, Overlain by Pliocene Alkali Basalt
Figure 2. Stratigraphic column for the McCoy Mountains Formation in its type section in the McCoy Mountains (Barth et al., 2004). Symbols to the right of the column show sample locations and U-Pb zircon minimum detrital ages.
Middle to upper crustal portions of two juxtaposed Mesozoic magmatic arcs are exposed in the area being investigated in this field trip, one being the Mojave province of the Cordilleran orogen and the other being the “exotic” or “suspect” Tujunga terrane (also termed the “San Gabriel Terrane”), that, in part, comprises the Transverse Ranges, including the San Gabriel Mountains, and which continues southeastward into the Mojave Desert to include desert mountain ranges south of exposures of the McCoy Mountains Formation, including the Chuckwalla, Little Chuckwalla, Chocolate, Pinto, Eagle, and Mule Mountains. The Tujunga terrane, a fault-bounded block underlying 21,000 km2 of southern California and adjacent Arizona, has been termed a suspect terrane due to its distinctive crystalline units that appear to have no correlative ties to native North America, including its exposure of 1190 Ma anorthosite (Barth et al., 2001a). The Tujunga terrane is everywhere allochthonous. Crystallization thermobarometry of dated plutons within the terrane has indicated that much of its apparent “suspect” nature stems from its partial derivation from the middle crust. Batholiths of the central and western Mojave have been studied by Coleman and Walker (1992), Coleman et al. (1992), and Miller and Glazner (1995). Large intrusions also comprise the central and eastern Mojave Desert and early studies include
those of Anderson and Rowley (1981), Beckerman et al. (1982), Miller et al. (1982), Howard et al., (1987), John (1987), Miller et al. (1990), Anderson and Cullers (1990), Miller et al. (1992), Anderson et al. (1992), Young et al., (1992), Miller and Wooden (1994), Gerber et al. (1995) and Mayo et al. (1998). The Mojave Desert region also contains other basement terranes of uncertain origin, including the Joshua Tree terrane. Bender et al. (1993) have concluded that the Joshua Tree terrane is contiguous with the Mojave and that the allochthonous Tujunga terrane represents a displaced, middle crustal section of the Mojave block. Mesozoic plutons of the Tujunga terrane (Barth et al., 1990; Barth et al., 1995) bear strong isotopic and elemental affinities to the plutons of the eastern Mojave, including essentially identical time-transgressive and compositional changes in magmatic arc construction, including both Early Proterozoic and Mesozoic plutons, thus suggesting that the Tujunga terrane is a displaced portion of the Mojave province (Bender et al., 1993; Anderson et al., 1992; Anderson et al., 1993). Mesozoic batholiths of the Mojave Desert and the Tujunga terrane occur in three distinct pulses (Fig. 3). Scattered Triassic intrusions, often distinctly K-feldspar megacrystic, were
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emplaced between 250 and 207 Ma (Barth and Wooden, 2006). After a magmatic lull of ca. 40 Ma, Jurassic magmatism led to widespread intrusion between 165 and 143 Ma. The Independence dike swarm was emplaced at 149 Ma late in this magmatic epoch (Carl and Glazner, 2002). Following another magmatic lull, one of 45 Ma, renewed magmatism in the mid Cretaceous progressed to a peak of activity at 82–72 Ma. Similar-aged magmatic episodes occur throughout the Cordilleran orogen of western North America, but nowhere else in the orogen is there the range of crustal depth of pluton emplacement as seen in the Mojave province, including age-correlative units in the Transverse Ranges. Mid-crustal plutons emplaced at depths greater than 20 km occur both in allochthonous sheets in Mesozoic thrust fault complexes and in the lower plates of Cenozoic metamorphic core complexes (Anderson, 1996). Interestingly, all Mesozoic plutons of the Mojave, regardless of age, were derived from high fO2 magmas, as shown by their high magnetite content. Low fO2 granitic magmas typically have ilmenite as the principle Fe-Ti oxide phase as found for considerable eastern portions of the Peninsular Ranges batholith (Shaw et al., 2003) Triassic Plutonism Several Triassic plutons occur in southern California including in the San Gabriel and Mule Mountains of the Tujunga terrane and the Granite I, Little San Bernardino and San Bernardino Mountains. U-Pb (zircon) age determinations range from 250 to 207 Ma (Barth et al., 1990; Miller, 1977; Frizzell et al., 1986). (Note: there are four named Granite Mountains in the Mojave Desert. What we term Granite I borders the San Bernardino Mountains and Granite II in the central Mojave Desert is adjacent to the Providence Mountains.) The plutons of both the Mojave province and the Tujunga terrane share many common megascopic and petrologic attributes making them much different from those of Jurassic and Cretaceous age. Metaluminous, high K2O + Na2O, and often of low silica (Fig. 4), the plutons are principally of monzonite and quartz monzonite with lesser amounts of monzodiorite and diorite. Most are
alkalic, with a shoshonitic affinity. High abundances of Ba, Sr (>1000 ppm) (Fig. 5), and light rare earth elements are also characteristic as first recognized by Miller (1977). Megacrystic K-feldspar is common. High feldspar content is reflected by low abundances of Fe, Mg, and Ti. The principal mafic minerals are hornblende and biotite ± clinopyroxene and garnet. The garnet coexisting with hornblende in the San Gabriel Mountains Mount Lowe intrusion (Tujunga terrane) and the quartz monzonite pluton of Granite I Mountains (Mojave province), is exceptionally enriched in grossular and andradite components. Barth (1990) has documented that the Mount Lowe was emplaced in the middle crust based on calculated pressures of 5.5–7.0 kbar, consistent with several mineralogical attributes of deep-seated crystallization, including markedly aluminous hornblende (to >11% Al203), calcic garnet, magmatic epidote, and siliceous primary muscovite. For the plutons in the San Bernardino and adjacent Granite I Mountains, Miller (1978) has argued that the magmas were derived from a large ion lithophile element-enriched, quartz eclogite source. The Mount Lowe intrusion is an immense, batholithic-sized (>300 km2), zoned plutonic complex that occurs in the San Gabriel Mountains and in correlative exposures east of San Andreas fault (Chocolate, Little Chuckwalla, Mule, and Trigo Mountains). Based on a broad database of elemental and Sr, Pb, and Nd isotopic data, Barth and Ehlig (1988) and Barth et al. (1990) have argued that the marginal zone of the intrusion was derived from low degrees of partial melting of an eclogitic and enriched subcontinental lithospheric source with virtually no input of continental material, whereas the central zone formed originally in a similar manner but with considerable crustal enrichment. Jurassic Plutonism Magmatic arc construction in the southern Cordillera and in the Tujunga terrane became a major feature of the orogen by the mid-Jurassic. The plutons are largely metaluminous, typically contain hornblende ± clinopyroxene and include gabbro, diorite, quartz monzodiorite, quartz monzonite, quartz syenite, and syenogranite. The granitic rocks are coarse grained and seriate to porphyritic with large, lavender-colored K-feldspar phenocrysts (Tosdal et al., 1989). The plutons have metasomatic effects not encountered in intrusions of Triassic or Cretaceous age. Zones of albitization occur in the Bristol, Ship, and Marble Mountains plutons. Based on stable isotopic data of the altered rocks, Fox and Miller (1990) have interpreted the fluids to be of meteoric origin. Replacement deposits of massive magnetite occur at upper crustal contacts with Paleozoic marbles in the Providence and Eagle Mountains, which have had considerable historical mining interest. Hall et al. (1988) present isotopic data supporting a magmatic origin of the fluids leading to these iron deposits. Crystallization thermobarometry shows a range of determined emplacement depth (Fig. 6). Upper crustal complexes occur in the Providence, Marble, Bristol, Chuckwalla, and Eagle Mountains (1–3 kbar). In contrast, deep-seated Jurassic complexes
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and 206Pb/204Pb defines a strikingly linear array, suggesting 1.6– 1.7 Ga crustal rocks were an important component of the source region (Anderson et al., 1990; Fox and Miller, 1990; J.L. Wooden, unpublished data). Young et al. (1992) modeled the origin of plutons in the Granite and Bristol Mountains by partial melting of hydrous and enriched mantle coupled with variable assimilativefractional crystallization involving ~10% Proterozoic crust.
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have been identified the Granite II and Cargo Muchacho Mountains (6–7 kbar) (Anderson, 1996, and references therein). Compositionally, these plutons are unlike those of Cretaceous or Triassic age but no significant difference occurs across the inferred position of the Mojave-Tujunga terrane boundary. Total alkalis are elevated, Fe/Mg ratios are intermediate (transitional between calc-alkaline and tholeiitic), and Sr abundances are moderate. Alkali-lime indexes require the existence of both calc-alkalic and alkali-calcic members. Limited isotopic data show δ18O values clustering near 8‰, Sri ranging from 0.706 to 0.709, and εNdt from −6 to −9 (Fig. 7). Covariation of 207Pb/204Pb
Cretaceous Plutonism After extended post-Jurassic magmatic lull, voluminous Cretaceous igneous activity in both the Tujunga terrane and the Mojave region began at ca. 95 Ma and reached a peak at 82– 72 Ma (Miller et al., 1982; Beckerman et al., 1982; Wright et al., 1986, 1987; Anderson et al., 1990). As found for the Jurassic plutons, the depths of emplacement of the Cretaceous vary widely (Anderson, 1996). Mid-crustal intrusions have been found in the San Gabriel, Old Woman, Granite II, Chemehuevi, and Whipple Mountains (Anderson et al., 1988). Plutons in Joshua Tree National Park were largely intruded at shallow crustal levels, but include a mid-crustal section (Needy et al., 2006); upper crustal intrusions have been identified in the Teutonia and Cadiz Valley batholiths, the Chuckwalla Mountains, and the Sacramento Mountains core complex. Compositionally, and by rock type, the Cretaceous suites of the Mojave and the Tujunga terrane are distinct from the magmatic activity of earlier Mesozoic intrusions. None are alkalic. The K2O abundance is lower (at medium to high K) and most plutons are relatively silicic (>66 wt% SiO2) and calc-alkaline. Two-mica granites are common, as are metaluminous hornblende-biotite-sphene granodiorites. Sr contents are usually less than 800 ppm, except for the calcic and Sr-rich granitoids of the Whipple core complex. Available isotopic data are currently limited to plutons in the Whipple, Chemehuevi, San Gabriel, and Old Woman Mountains and include δ18O values from 7 to 9‰,
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Sri from 0.706 to 0.711 (to 0.719 for the strongly peraluminous granites of the Old Woman Mountains), and εNdt from −10 to −17 (John and Wooden, 1990; Anderson and Cullers, 1990; Barth et al., 1995; Miller et al., 1990). Tested models show a broad range of magma origins for the above complexes, including significant derivation from heterogeneous Proterozoic crust, but with some variable input of radiogenic mantle and/or mafic crust mixed with a Proterozoic component. The main compositional difference between the Cretaceous plutons is seemingly related to depth of emplacement. While the shallowly emplaced (100 m.y.), they ALWAYS form sheeted bodies, suggesting that emplacement style is consistent over time; (5) they have a dramatic range in composition (gabbros to granites and cumulates to late aplites and pegmatites) over spatially short distances that are not organized into patterns, such as in normally or reversely zoned plutons; and (6) sheeted plutons are usually strongly deformed in the magmatic or solid state indicating the presence of long-lived or recurrent deviatoric stress in the mid crust.
Eastern Transverse Ranges and southern Mojave Desert For example, if we consider Mesozoic magmatism in the North American Cordillera the following prominent sheeted complexes have been recognized: (1) The “Great Tonalite Sill” (really a collection of sills or dikes) in the Coast Plutonic Complex; (2) the Skagit Gneiss Complex in the Cascades crystalline core, Washington; (3) sheeted complexes in the Idaho suture zone; (4) a central sheeted complex in Peninsular Ranges Batholith that extends from southern California to at least as far south as the Sierra San Pedro Martir; and (5) the sheeted complex in the Keys View quadrangle. Such magmatic sheeted zones have been interpreted in a number of ways, such as: (1) the deeper parts of magma plumbing systems that are feeding more elliptically shaped chambers at shallower levels (e.g., Cascades core); (2) magmatism controlled by emplacement into an active fault or faults (e.g., Great Tonalite Sill); (3) magmatism along fundamental lithospheric-scale boundaries that may or may not be faults (e.g., Idaho suture zone and Peninsular Ranges Batholith examples, interpreted to represent a transition from continental to oceanic crust, with or without major faults); (4) dike swarms controlled by a regional stress field; (5) sheeted bodies (whether or not dikes) in which emplacement is strongly controlled by preexisting host rock anisotropy; (6) zones of syn-emplacement extension; and (7) combinations of the above. Directions to Stop 7 Depart camp, travel southeast on Quail Springs Road to the intersection with Keys View Road. Continue south on Keys View Road to Stop 7 [0577200 3760650]. Stop 7. Hike West to View Late Cretaceous Palms Granite and Intrusive Contact with Paleoproterozoic Gneiss The grain size, apparent composition, and weak fabrics in the Late Cretaceous Palms pluton do not change their general appearance as the margin is approached, although local variations occur. The intrusive contact is often quite discordant to the metamorphic layering and dominant foliation in the gneisses, but there are some local zones where older structures are rotated into subparallelism with the Palms contact. Some xenoliths (both rafts and stoped blocks) certainly occur along the margin. The intrusive contact changes its orientation from place to place but the overall pattern is that it dips fairly steeply to the northeast. There is one interesting area where the Blue granodiorite has a gently dipping upper intrusive contact with the gneisses, which is then truncated by the fairly steeply dipping Palms contact. This is one location that would indicate that the Palms did intrude more gently dipping sheeted intrusive rocks and gneisses.
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Stop 8. Hike East to View Proterozoic Gneiss and Contact with Late Cretaceous Blue Granodiorite In the Keys View quadrangle (Fig. 10), we can examine a wonderful transition between fairly circular (in map view) plutons that in the field appear fairly homogeneous and weakly foliated (Palms granite, Squaw Tank granite), to more elongate plutons with more intense fabrics and clear evidence for multiple internal sheets (e.g., Blue granodiorite, Stubbe Springs granite), to a very heterogeneous sheeted complex showing a range of compositions from gabbro to two mica granites, both Cretaceous and Jurassic ages, and with variable fabric intensities (Quail Mountain complex, Bighorn complex). At this stop, we will hike through the transition from Paleoproterozoic gneisses into the internally sheeted Blue granodiorite. Gneisses here are a mixture of both orthogneisses and paragneisses with local amphibolite. All are intensely deformed—we see clear evidence for early isoclinal folds with the dominant fabric parallel to their axial planes. These early isoclines fold a previous foliation and metamorphic layering and are in turn refolded by more open, upright folds. The regional metamorphic grade was at least amphibolite facies. We also observe local development of andalusite, garnet, and sillimanite near the contact with the Blue granodiorite, but more work is needed to sort out the differences between Paleoproterozoic regional and Mesozoic contact metamorphism. Structures in the gneisses are discordantly intruded along a fairly steeply dipping contact of the Blue granodiorite, although the usual local zones of subparallelism exist. The intrusive contact takes a number of sharp steps and we do find xenoliths (some large enough to be mapped) in the granodiorite. Along this eastern margin, the Blue granodiorite tends to be more homogeneous (less sheeted) and has numerous mafic clots (small enclaves?) in it. Magmatic fabrics steepen near the margin but are typically fairly gently southwest dipping (foliation) or gently northwestsoutheast–plunging (lineation) in this body. Small compositional and textural changes define internal sheets in the granodiorite. More dramatic sheeting defined by 2-mica, garnet granites also occurs. Small pods of Stubbe Springs–like granites intrude the granodiorite, particularly as its western margin is approached. Directions to Stop 9 Depart stop, travel west to return to Keys View Road, continue south on Keys View Road to the Keys View parking area, Stop 9 [0575226 3754200]. Stop 9. Hike to View Sheeted Granodiorite and Granite along Western Contact of Late Cretaceous Blue Granodiorite
Directions to Stop 8 Depart stop, continue south on Keys View Road to Lost Horse Mine (dirt) Road, turn left and continue southeast on Lost Horse Mine Road to parking area [0576530 3757780].
The western margin of the Blue granodiorite is often much more cryptic because of the sheets within the granodiorite and younger intrusions along this margin. We will hike along a transect perpendicular to this margin to examine the granodiorite
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and a number of granitic bodies intruding into the granodiorite and end near Keys View where the midcrustal Bighorn complex is exposed. Here the Blue granodiorite is a medium-grained, hornblende biotite granodiorite with a moderately well developed northweststriking, northeast-dipping magmatic foliation and northwest trending lineation. Regionally, the magmatic foliation defines an open synformal fold pattern with the northwest-southeast–trending fold axis subparallel to the mineral lineation. Locally, zones of subsolidus deformation define steep zones of southwest-sideup shear. Bodies of 2-mica garnet granite intrude the granodiorite. These bodies are less strongly foliated than the granodiorite and have shapes ranging from sheet-like to blobby. Farther southwest, granodiorites are increasingly intruded by a complex variety of sheet-shaped bodies, locally separated by screens of orthogneiss and paragneiss, and varying in composition from gabbros and hornblende cumulates to 2-mica granites. Sheets vary in thickness from less than a meter to ≥10 m and generally dip moderately to the northeast. Intrusive relationships are complex, but there is a general pattern of more mafic being intruded by more
felsic sheets. Many sheets have well-developed magmatic fabrics locally overprinted by weak subsolidus deformation. These magmatic fabrics are more intensely developed in granodiorites and are typically weaker in granites (Brown et al., 2006). End of Day 2 DAY 3 In Days 1 and 2 of this trip, we traveled from east to west through progressively deeper levels of the Cordilleran magmatic arc and its Proterozoic wall rocks. This morning we will visit the northwestern Orocopia Mountains, which in their higher structural levels include crustal rocks of North American affinity similar to those viewed previously. However, the Orocopia Mountains also provide a window into an oceanic complex which was underplated beneath much of southern California and adjoining regions during the Late Cretaceous to early Cenozoic Laramide orogeny. Here known as the Orocopia Schist, these rocks are part of the larger Pelona-Orocopia-Rand Schist (Fig. 11; Haxel et al., 2002; Jacobson, et al., 2007). Emplacement of the schists beneath
Figure 11. Geology of the Orocopia Mountains and vicinity (after Crowell, 1975). Inset shows distribution of Pelona-Orocopia-Rand Schists. Gf—Garlock fault, LA—Los Angeles, SAf—San Andreas fault.
Eastern Transverse Ranges and southern Mojave Desert the arc terrane represents a first order tectonic event that involved the stripping away of the lowermost North American crust and the entire thickness of underlying mantle lithosphere. The original subduction thrust responsible for this event is preserved in only a few areas. In most cases it has been replaced by syn-subduction (i.e., Laramide age) and/or middle Cenozoic low-angle normal faults. In the Orocopia Mountains, there is indirect evidence for Laramide exhumation, but the primary contact between schist and arc terrane (here referred to as the “upper plate”) is a major Miocene detachment fault (Orocopia Mountains detachment fault; Robinson and Frost, 1996; Ebert, 2004; Jacobson et al., 2007). Below we provide brief descriptions of the schist and upper plate. Orocopia Schist The schist in the Orocopia Mountains is composed dominantly of metagraywacke interpreted as trench sediment (Grove et al., 2003; Jacobson et al., 2007). Detrital zircons indicate a depositional age no older than ca. 70 Ma and perhaps as young as 62 Ma (Fig. 12; but note that the schist protolith in other
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areas is as old as 90 Ma). The schist also includes various rock types inferred to have been scraped off the Farallon plate. These include a few percent of metabasite with normal to enriched mid-ocean ridge basalt composition and even lesser amounts of Fe-Mn metachert, marble, serpentinite, and talc-actinolite rock (Haxel et al., 2002). Hornblende 40Ar/39Ar ages of 54–50 Ma and muscovite ages as old as ca. 50 Ma (Fig. 12) demonstrate that underplating beneath North American crust and initial cooling and exhumation occurred shortly after deposition of the graywacke protolith in the trench. Biotite ages of 30–20 Ma and Kfeldspar multi-diffusion domain analysis reveal a second, very rapid, period of exhumation at ca. 24–22 Ma (Fig. 12). The latter is attributed to normal-sense slip on the Orocopia Mountains detachment fault. In most of the range, prograde assemblages in the schist belong to the albite-epidote amphibolite faces. However, the upper half of the exposed section shows extensive retrogression to greenschist facies. We believe that this overprint is related to the first (early Cenozoic) phase of exhumation. A thin (several m to several 10s of m) mylonite zone right at the top of the schist is
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considered to be a younger feature associated with the Orocopia Mountains detachment fault. Upper Plate The upper plate includes three main lithologic units: Proterozoic gneiss, 1.2 Ga anorthosite-syenite, and 76 Ma leucogranite (Crowell, 1962, 1975; Crowell and Walker, 1962; Barth et al., 2001a; Jacobson et al., 2007). The gneiss is locally intruded by the anorthosite-syenite, but in general these two units are spatially distinct. In contrast, the gneiss is intimately intruded by the leucogranite, so we combine these two rock types in a single map unit (leucogranite-gneiss) despite their disparate ages. In most areas, the anorthosite-syenite unit sits structurally above the leucogranite-gneiss unit along a low-angle fault referred to as the upper plate detachment fault (Figs. 11 and 13). Additional lowangle faults may also be present within the anorthosite-syenite and leucogranite-gneiss units. The anorthosite-syenite unit includes anorthosite, gabbro, and syenite, as well as compositions (mangerite and jotunite) intermediate between these three end members. Retrograde metamorphism of inferred Proterozoic age is indicated by essentially universal replacement of primary igneous pyroxene by hornblende and biotite, presence of hairline mesoperthite, and local development of blue quartz in the more felsic units. Structurally lower levels of the anorthosite-syenite unit are intruded by dikes and small stocks of the same leucogranite that abundantly intrudes the gneiss. The gneiss is largely quartzofeldspathic but also includes a mafic component. Metamorphism was largely to amphibolite faces, but reached granulite facies near intrusive contacts with the anorthosite-syenite. As in the anorthosite-syenite, most granulite assemblages have been retrograded to amphibolite facies, which are commonly associated with blue quartz. The leucogranite associated with the gneiss generally has an aplitic texture. It consists mostly of subequal amounts of quartz, plagioclase, and K-Feldspar with minor biotite that is commonly retrograded to chlorite. The various structural levels of the upper plate show diverse cooling histories (Fig. 12). In all cases, however, the early to middle Cenozoic history is substantially different from that of the Orocopia Schist. It is this disparity which indicates that the Orocopia Mountains detachment fault can be no older than early Miocene. Directions to Stop 10 Depart camp, head southeast on Pinto Basin Road to return to south park entrance station, continue south on Cottonwood Pass Road to southern park boundary and intersection with I-10. Cross I-10 and continue heading south. In 0.4 mi [611074 3724293], turn left onto poorly maintained paved road. In 0.05 mi [611150 3724261], turn right onto dirt road that heads SSW up the alluvial fan toward the Orocopia Mountains. In 0.75 mi, pass power line. In another 0.65 mi, the road enters a canyon in modest hills of
Figure 13. Stop locations, day 3. See Figure 11 for location and explanation of fault hachures. UTM 1 km grid ticks indicated. Abbreviations: an-sy—anorthositesyenite unit, lg-gn—leucogranite-gneiss unit, os—Orocopia Schist.
porphyritic granodiorite. Our limited dating indicates a possible Jurassic age. The road skirts the west side of the canyon. Approximately 0.4 mi after entering the canyon [610368 3721575] take the road that leaves the main wash heading SSW. In 0.5 mi [610174 3720855] pass a small hill of Diligencia Formation. This unit sits unconformably upon the granodiorite we just passed in the canyon. The road approaches a major drainage. At 0.3 mi past the outcrop of Diligencia Formation [610279 3720381] veer to the right off the road traversing the alluvial fan and descend into the wash (it is easy to miss this intersection). Drive up the wash. Pass highly altered outcrops of leucogranite in the low walls of the wash. These are located on the northeast side of the Clemens Well fault, which may be a strike-slip fault with modest to large displacement or a steepened detachment fault related to the Orocopia Mountains detachment fault and upper plate detachment fault (Crowell, 1962, 1975; Powell, 1993; Robinson and Frost, 1996). Continuing up the wash, we cross the Clemens Well fault and pass into the leucogranite-gneiss unit. Stop 10 [610556 3718428]. Start of Traverse, Upper Plate Rocks This is as far as one can drive up the wash (~1.5 mi upstream from our entry point into the wash). Vehicle maneuverability is limited and for the field trip we may park a bit below this point. From here we will take a moderate hike up the canyon examining the leucogranite-gneiss and anorthosite-syenite units of the upper plate and the underlying Orocopia Schist. Particular emphasis will be placed on the early Miocene fault contacts between these units. If time permits, we will climb up to the main ridge of the Orocopia Mountains for a view into the Salton Trough. We will examine rocks all along the traverse, with a few key localities broken out as stops.
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We start our hike in the leucogranite-gneiss unit. Note the high degree of brittle deformation and hydrothermal alteration. The leucogranite cuts most of the ductile fabric in the upper plate, but locally shows a modest foliation itself.
Return to vehicles. Depart stop, return to I-10, east on I-10 to Blythe, north on U.S. Highway 95 to Las Vegas.
Stop 11 [610992 3717799]. Upper Plate Detachment Fault
The National Science Foundation (EAR-0106881, 0408730 and 0711119 to APB, EAR-9902788 and EAR-0106123 to CEJ), the National Geographic Society (7214-02), the National Park Service, the Southern California Earthquake Center, Department of Energy, and the Joshua Tree National Park Association provided support for this research. We are grateful to the staff at Joshua Tree National Park for supporting our work in the park, and to Kenneth Brown, Kristin Ebert, Nicole Fohey, Marty Grove, Kristin Hughes, Vali Memeti, Sarah Needy, Emerson Palmer, Jane Pedrick, Geoff Pignotta, Kelly Probst, and Ana Vućić who collaborated with us in the research summarized here. Robert Powell, Paul Stone, Richard Tosdal, and Dee Trent provided guidance and many helpful discussions, and Calvin Miller and Gene Smith provided constructive reviews of the manuscript.
This is a truly outstanding exposure of the contact between the leucogranite-gneiss and anorthosite-syenite units (upper plate detachment fault). The leucogranite-gneiss is exposed at the base of the outcrop. It is overlain along a sharp, low-angle contact by the basal part of the anorthosite-syenite unit which itself is cut by highly sheared dikes of leucogranite. This middle slice is in turn overlain along another low-angle fault by a structurally higher level of the anorthosite-syenite unit without leucogranite. Note the well developed gouge and other indicators of brittle deformation along the various contacts, demonstrating the relatively shallow nature of this fault system. Sense-of-shear indicators are not common but generally show top-NE to top-E sense of movement. Note the steep faults of various orientations within the different structural plates. Stop 12 [610947 3717648]. Orocopia Mafic Schists Mafic Orocopia Schist with albite porphyroblasts up to 4– 5 mm in diameter. Other major prograde minerals are hornblende and epidote (albite-epidote amphibolite facies). Secondary chlorite is widespread in thin section. Metagraywacke in adjacent outcrops is dominated by quartz, albite, and muscovite with lesser biotite and garnet. Secondary chlorite is present in both units. Stop 13 [611070 3717546]. Orocopia Mountains Detachment Fault Contact between Orocopia Schist and the part of the anorthosite-syenite unit intruded by leucogranite dikes (as seen in the middle plate of Stop 11). Note the mylonite and asymmetric shear bands in the schist, but brittle deformation only in the anorthosite-syenite unit. This contrast in structural behavior confirms the thermochronologic evidence (Fig. 12) that this contact is a normal fault. Stop 14 [611325 3717297]. Orocopia Mountains Detachment Fault The contrast between brittle deformation within the upper plate but mylonitization of the schist is similar to that seen at Stop 13. Here the upper plate is transitional in composition between the leucogranite-gneiss and anorthosite-syenite units (i.e., it includes rocks typical of the anorthosite-syenite suite but exhibits an exceptionally high degree of intrusion by the leucogranite). If time permits, we will walk up this contact to the main Orocopia ridge. Note the mylonitic nature of the schist along the contact, similar to the textures observed at Stop 14.
ACKNOWLEDGMENTS
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Printed in the USA
The Geological Society of America Field Guide 11 2008
Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary, northwest Arizona: A tale of three basins, immense lacustrine-evaporite deposits, and the nascent Colorado River James E. Faulds* Nevada Bureau of Mines and Geology, MS 178, University of Nevada, Reno, Nevada 89557, USA Keith A. Howard* U.S. Geological Survey, MS 973, Menlo Park, California 94025, USA Ernest M. Duebendorfer* Department of Geology, Northern Arizona University, Flagstaff, Arizona 86011, USA ABSTRACT In northwest Arizona, the relatively unextended Colorado Plateau gives way abruptly to the highly extended Colorado River extensional corridor within the Basin and Range province along a system of major west-dipping normal faults, including the Grand Wash fault zone and South Virgin–White Hills detachment fault. Large growth-fault basins developed in the hanging walls of these faults. Lowering of base level in the corridor facilitated development of the Colorado River and Grand Canyon. This trip explores stratigraphic constraints on the timing of deformation and paleogeographic evolution of the region. Highlights include growth-fault relations that constrain the timing of structural demarcation between the Colorado Plateau and Basin and Range, major fault zones, synextensional megabreccia deposits, nonmarine carbonate and halite deposits that immediately predate arrival of the Colorado River, and a basalt flow interbedded with Colorado River sediments. Structural and stratigraphic relations indicate that the current physiography of the Colorado Plateau–Basin and Range boundary in northwest Arizona began developing ca. 16 Ma, was essentially established by 13 Ma, and has changed little since ca. 8 Ma. The antiquity and abruptness of this boundary, as well as the stratigraphic record, suggest significant headward erosion into the high-standing plateau in middle Miocene time. Thick late Miocene evaporite and lacustrine deposits indicate that a long period of internal drainage followed the onset of extension. The widespread distribution of such deposits may signify, however, a large influx of surface waters and/or groundwater from the Colorado Plateau possibly from a precursor to the Colorado River. Stratigraphic relations bracket arrival of a through-flowing Colorado River between 5.6 and 4.4 Ma. Keywords: Basin and Range, Colorado River, extension, paleogeography, Colorado Plateau *
[email protected];
[email protected];
[email protected] Faulds, J.E., Howard, K.A., Duebendorfer, E.M., 2008, Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary, northwest Arizona: A tale of three basins, immense lacustrine-evaporite deposits, and the nascent Colorado River, in Duebendorfer, E.M., and Smith, E.I., eds., Field Guide to Plutons, Volcanoes, Faults, Reefs, Dinosaurs, and Possible Glaciation in Selected Areas of Arizona, California, and Nevada: Geological Society of America Field Guide 11, p. 119–151, doi: 10.1130/2008.fld011(06). For permission to copy, contact
[email protected]. ©2008 The Geological Society of America. All rights reserved.
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INTRODUCTION
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In northwest Arizona, the Colorado River crosses an unusually abrupt boundary between the Colorado Plateau and the Basin and Range province (Fig. 1). Essentially flat, relatively unextended strata on the high-standing Colorado Plateau give way to moderately to steeply tilted fault blocks in the Basin and Range province across a system of west-dipping normal faults that includes the Grand Wash fault zone (Lucchitta, 1966, 1979) and South Virgin–White Hills detachment fault (Fig. 2). Unlike other parts of the Colorado Plateau–Basin and Range boundary (e.g., southwest Utah and central Arizona), a broad transition zone is missing in northwest Arizona (Fig. 1). Instead, a 100km-wide region of highly extended crust within the Basin and Range, referred to as the northern Colorado River extensional corridor (Faulds et al., 1990), directly borders the Colorado Plateau on the west. Within the footwall of the Grand Wash fault zone, the western edge of the Colorado Plateau is marked by
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Figure 1. Digital elevation model showing the abrupt western margin of the Colorado Plateau. In contrast to broad transition zones throughout much of Utah and Arizona, the Colorado Plateau gives way abruptly westward to the Basin and Range province in the Lake Mead region of northwestern Arizona. Small white box encompasses the study area in the southern White Hills. The map projection is cylindrical and equidistant with the shape corrected for 37.5° north latitude. Lighting is from the northwest.
the imposing, west-facing fault-line escarpment of the Grand Wash Cliffs, which consist of subhorizontal Paleozoic strata rising ~1.3 km above several east-tilted half grabens in the corridor, including the Grand Wash trough and the Hualapai basin. With respect to the base of the Tertiary section, structural relief across the Grand Wash fault zone commonly exceeds 5 km. Approximately 15–30 km west of the Grand Wash Cliffs, the gently west-dipping South Virgin–White Hills detachment fault dissects the corridor. The South Virgin–White Hills detachment fault is one of the most prominent structures in the northern part of the corridor, as it accommodated as much as 17 km of normal displacement and has many characteristics of classic detachment faults (Duebendorfer and Sharp, 1998; Brady et al., 2000). Thus, the transition between the essentially unextended Colorado Plateau to the highly attenuated Basin and Range occurs across a relatively narrow ~30-km-wide region in northwest Arizona. It is noteworthy that the Colorado River flows transversely across this abrupt strain gradient, having excavated the Grand Canyon within the western part of the Colorado Plateau and traversing orthogonal to the structural grain within the Lake Mead region in the northern part of the extensional corridor (Figs. 2 and 3). The evolution of the Colorado River and Grand Canyon have long fascinated geoscientists, and many models have been proposed for its development (e.g., Powell, 1875, 1895; Blackwelder, 1934; Longwell, 1946; Hunt, 1969; Lucchitta. 1966, 1972, 1979, 1989; Young and Spamer, 2001). Recent work has greatly refined the evolution of the Colorado River, particularly the timing of inception for reaches downstream of the Grand Canyon (Spencer et al., 2001; Faulds et al., 2001b, 2002a; House et al., 2005; Dorsey et al., 2007), models for drainage development (Spencer and Pearthree, 2001; House et al., 2005), and rates of incision within the Grand Canyon (Fenton et al., 2001; Pederson and Karlstrom, 2001; Pederson et al., 2002). However, the relationships between major precursor events and development of the Grand Canyon and lower Colorado River have received less attention. Clearly, the structural and topographic foundering of the Basin and Range province, particularly within the Colorado River extensional corridor, promoted excavation of at least the western part of the Grand Canyon within the high-standing Colorado Plateau. Understanding the spatial and temporal patterns of deformation within the extensional corridor is therefore critical for establishing a physiographic, structural, and temporal framework by which to assess the evolution of the Colorado River and associated drainage systems. On this field trip, we will evaluate the timing and nature of Cenozoic structural demarcation between the Colorado Plateau and the Basin and Range province in northwest Arizona (Fig. 3), as chronicled in the stratigraphy of major half grabens in the hanging walls of the Grand Wash and South Virgin–White Hills fault zones. A major goal of the trip is to further elucidate the relations between the stratigraphy and deformational history of these basins with the evolution of the Colorado River.
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Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary
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Late Tertiary-Quaternary basin fill sediments Late Miocene basalts at SP and TM and Pliocene basalts elsewhere (only larger exposures shown) E-tilted Miocene volcanic and sedimentary strata W-tilted Miocene volcanic and sedimentary strata Miocene plutons and dike swarms Mainly Paleocene to early Miocene sedimentary rocks on Colorado Plateau
Gently dipping normal fault (dashed where concealed) Moderately to steeply dipping normal fault (ball on downthrown side) Strike-slip fault (arrows show relative sense of movement Thrust fault Axial part of Black Mountains accommodation zone, dashed where concealed, showing anticlinal and synclinal segments
Late Cretaceous plutons Anticline Paleozoic and Mesozoic sedimentary strata Proterozoic plutonic and metamorphic rock
Syncline
Figure 2. Generalized geologic map of the northern Colorado River extensional corridor. The box surrounds area covered by the field trip. Basins: DB—Detrital basin; GB—Gregg basin; GT—Grand Wash trough; HB—Hualapai basin; NWH—northern White Hills basin; LVV—Las Vegas Valley; OA—Overton Arm; SAB—Sacramento basin; SWH—southern White Hills basin. Faults: BG—Blind Goddess fault; CMF— Cerbat Mountains fault; CY—Cyclopic fault; DF—Detrital fault; LBF—Lost Basin Range fault; LMF—Lakeside Mine fault; LMFS—Lake Mead fault system; LVVSZ—Las Vegas Valley shear zone; MS—Mountain Spring fault; NGW—northern Grand Wash fault; SGW—southern Grand Wash fault; SIF—Saddle Island fault; SSW—Salt Spring Wash fault; SVWHD—South Virgin-White Hills detachment fault; WHF— White Hills fault; WF—Wheeler Ridge fault. Major physiographic features: BRP—Basin and Range province; CB—central Black Mountains; CL—Callville Mesa; CP—Colorado Plateau; CM—Cerbat Mountains; DS—Dolan Springs; FM—Frenchman Mountain; GC—Grand Canyon; GM—Garnet Mountain; GWC—Grand Wash Cliffs; HM—Hiller Mountains; HR—Highland Range; LB—Lost Basin Range; LM—Lake Mead; ML—Lake Mohave; MM—Muddy Mountains; MV—Meadview; NB—northern Black Mountains; NE—northern Eldorado Mountains and basin; NM—Newberry Mountains; SB—southern Black Mountains; SE—southern Eldorado Mountains; SP—Snap Point; TB—Temple Bar; TM—Table Mountain Plateau; WHR—Wheeler Ridge. Other structures: AZ—Black Mountains accommodation zone; GBB—Gold Butte block; MPB—Mount Perkins block.
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Figure 3. Map showing field trip route for each day and major physiographic features in the northern Colorado River extensional corridor and neighboring parts of the western margin of the Colorado Plateau.
REGIONAL GEOLOGIC SETTING Northwestern Arizona and southern Nevada have occupied a critical position in the western Cordillera since Mesozoic time. In Cretaceous to early Tertiary time, this region marked the northern edge of a large crystalline terrane referred to as the Kingman uplift (Goetz et al., 1975) or Kingman arch (Bohannon, 1984), which was stripped of its Paleozoic and Mesozoic cover by erosion during Paleogene time. Later, during the middle to late Tertiary, the Lake Mead region marked the northern end of the highly extended Colorado River extensional corridor, which
was dominated by ~east-west extension (e.g., Davis et al., 1980; Howard and John, 1987; Davis and Lister 1988; Faulds et al., 1990, 2001a; Campbell and John, 1996). In contrast, the region directly north of Lake Mead was characterized by a complex three-dimensional strain field involving strike-slip faulting and north-south shortening, in addition to large-magnitude ~east-west extension (e.g., Weber and Smith, 1987; Anderson and Barnhard, 1993; Anderson et al., 1994; Duebendorfer and Simpson, 1994). In the northern part of the extensional corridor, calc-alkaline magmatism and major east-west extension swept northward in early to middle Miocene time (e.g., Glazner and Bartley,
Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary 1984; Gans et al., 1989; Faulds et al., 1999). Magmatism began ca. 18–20 Ma, 1–4 m.y. before the onset of major east-west extension (Faulds et al., 1995, 1999, 2002b; Gans and Bohrson, 1998). Voluminous, early Miocene, generally intermediate composition magmatism was accompanied by little deformation, although mild north-south extension affected some areas (Faulds et al., 2001a). Major east-west extension then battered the region beginning ca. 16–17 Ma in the south and migrating north-northwest to the western Lake Mead region by ca. 13 Ma. Extension ended in most areas by 11–8 Ma. Tertiary extension was accommodated by mainly west-dipping normal faults and east tilting of fault blocks in the Lake Mead area (e.g., Anderson, 1971, 1978; Duebendorfer and Sharp, 1998), whereas east-dipping faults and west-tilted fault blocks dominated to the south in the Lake Mohave region (e.g., Faulds et al., 1995; Fig. 2). The boundary between these oppositely dipping normal fault systems has been referred to as the Black Mountains accommodation zone, which corresponds to a 5–10-km-wide region of intermeshing, oppositely dipping normal faults and abundant extensional folds (Faulds et al., 1990, 2001a, 2002b; Faulds and Varga, 1998; Varga et al., 2004). The east- and west-tilted domains on either side of the accommodation zone are termed the Lake Mead and Whipple domains, respectively (Spencer and Reynolds, 1989). Estimates of extension within the northern part of the corridor range from ~75%–100% (e.g., Faulds et al., 1990; Brady et al., 2000). Thick sections (generally >3 km) of Tertiary volcanic and sedimentary strata rest directly on Proterozoic and late Cretaceous metamorphic and plutonic rock within the bulk of the extensional corridor (Anderson, 1971; Sherrod and Nielson, 1993; Faulds et al., 1995, 2002b; Beard, 1996). Sections are thickest in middle to late Miocene half grabens. The strata typically range in age from early to late Miocene and consist of mafic to felsic lavas, ash-flow tuffs, clastic sedimentary rocks, rock avalanche deposits, volcanic breccia, and evaporites. Although preserved to the north, east, and west of the region, Paleozoic and Mesozoic strata are missing from all but the northernmost part of the extensional corridor (i.e., Lake Mead region) owing to significant early Tertiary erosion of the Kingman arch. Basement rocks include Paleoproterozoic gneisses, ca. 1.4 Ga granite, Late Cretaceous–early Tertiary peraluminous (two-mica and garnet-bearing) granites, and early to middle Miocene silicic to intermediate plutons and mafic to felsic dike swarms. During extension, older units were progressively tilted to steeper dips concurrent with deposition of younger sequences on subhorizontal surfaces. Consequently, many of the basins contain well-developed tilt fanning (i.e., growth-fault sequences), whereby tilts within the synextensional parts of the section progressively decrease upwards. Volcanic units in many half grabens permit precise dating of the timing of extension (e.g., Faulds et al., 1995, 1999, 2002b; Duebendorfer and Sharp, 1998; Gans and Bohrson, 1998; Varga et al., 2004). Tilt fanning indicates that major eastwest extension began 16.7–15.7 Ma and continued at high rates until ca. 13 Ma in a broad region of the corridor extending from the latitude of Kingman, Arizona, on the south to the eastern Lake
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Mead region on the north (Anderson et al., 1972; Beard, 1996; Faulds et al., 1995, 1999, 2002b; Duebendorfer and Sharp, 1998). Major extension then shifted northwestward ca. 13 Ma into the western Lake Mead region, where it continued until ca. 9 Ma (Duebendorfer and Wallin, 1991; Harlan et al., 1998; Castor et al., 2000). Since ca. 8 Ma, the northern Colorado River extensional corridor has experienced only minor tilting and faulting. In contrast to the extensional corridor, the Colorado Plateau has remained tectonically stable through Cenozoic time and is essentially unextended at upper-crustal levels, as evidenced by subhorizontal Paleozoic and Mesozoic strata. Approximately 2 km of Paleozoic and Mesozoic strata caps the Colorado Plateau but is absent in much of the Basin and Range province of central and western Arizona (Peirce, 1985; Lucchitta and Young, 1986). Late Cretaceous marine deposits on the Colorado Plateau (Nations, 1989) suggest nearly 2 km of uplift during Cenozoic time (Parsons and McCarthy, 1995). However, the timing and nature of Colorado Plateau uplift remain controversial, because the transition zone and much of the Basin and Range province are structurally higher than the Colorado Plateau. Thus, both the Basin and Range and Colorado Plateau may have originally been uplifted, perhaps in early Tertiary time, but parts of the Basin and Range province later subsided. The border between the Colorado Plateau and Basin and Range is generally marked by a broad (~50–150 km wide) transition zone containing characteristics of both provinces (Peirce, 1985). In northwestern Arizona, however, an abrupt boundary separates the relatively unextended Colorado Plateau and Basin and Range province (Figs. 1 and 2). Large Miocene half grabens along the eastern margin of the Colorado River extensional corridor chronicle the evolution of this tectonic boundary and also elucidate major events that facilitated development of the Colorado River. A TALE OF THREE BASINS In this section, we describe three major basins along the eastern margin of the Colorado River extensional corridor: the Grand Wash trough, White Hills basin, and Hualapai basin. The Grand Wash trough and White Hills basin are complex, composite easttilted half grabens, whereas the Hualapai basin is a relatively simple east-tilted half graben. Neogene deposits within each of these half grabens have important implications for understanding the tectonic and paleogeographic evolution of this region. 40Ar/39Ar geochronology and geochemical correlations of tephras (tephrochronology) constrain the timing of deformation in both the Grand Wash trough and White Hills basin. The field trip will visit the Grand Wash trough on Day 1, northern White Hills basin on Day 2, and southern White Hills and Hualapai basins on Day 3. Grand Wash Trough (Day 1) The Grand Wash trough consists of at least two east-tilted half grabens, which are separated by Wheeler Ridge in the north and the Lost Basin Range in the south (Fig. 2). The eastern half
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Figure 4. View looking north at the Grand Wash trough, Grand Wash Cliffs, and Wheeler Ridge from Airport Point at the north end of Grapevine Mesa. Note the contrast between the subhorizontal strata along the Grand Wash Cliffs and steeply east-dipping strata on Wheeler Ridge.
Colorado Plateau
GC
GWC
Figure 5. View looking east of the Grand Wash trough and western margin of the Colorado Plateau, including Airport Point (AP), Grand Canyon (GC), Grand Wash Cliffs (GWC), Grapevine Mesa (GM), which is capped by the Hualapai Limestone), Grapevine Canyon (GrC), Sandy Point (SP), and Wheeler Ridge (WR). Grapevine Mesa essentially marks the floor of a late Miocene lake that immediately predates arrival of the Colorado River. Also, note the gentle northeast dip of Paleozoic strata along the Grand Wash Cliffs.
GrC
Grand Wash Trough
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Lake Mead
Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary graben developed in the hanging wall of the west-dipping northern Grand Wash fault and is centered in the Grapevine Wash area. To the west, the Gregg Basin is a relatively narrow east-tilted half graben that lies in the hanging wall of the west-dipping Wheeler Ridge and Lost Basin Range faults. The Wheeler Ridge and Lost Basin Range faults probably represent splays of the Grand Wash fault zone. As Wheeler Ridge dies out to the north of Lake Mead, the two half grabens coalesce to form a large composite basin, at least at exposed levels. Dissection by the Colorado River and its tributaries has produced excellent exposures of the upper part of the Tertiary section in both the Grand Wash trough and Gregg Basin (Figs. 4 and 5). The middle to late Miocene section within the Grand Wash trough (referred to as the rocks of the Grand Wash trough after Bohannon, 1984) includes, in ascending order, at least 250 m of middle to late Miocene fanglomerate, more than 120 m of a sandstone-siltstone facies with locally interbedded gypsum, and as much as 300 m of late Miocene limestone (Figs. 6 and 7A; Longwell, 1936; Lucchitta, 1966; Bohannon, 1984; Wallace,
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1999; Wallace et al., 2005; Blythe, 2005). The units interfinger and thicken eastward toward the deeper parts of the half graben. The conglomerate contains many large boulders (>5 m long) of the 1.4 Ga Gold Butte Granite (e.g., Volborth, 1962; Silver et al., 1977), a megacrystic rapakivi granite derived from the Gold Butte block in the south Virgin Mountains ~6–10 km to the west (Longwell, 1936; Lucchitta, 1966; Lucchitta and Young, 1986; Wallace, 1999; Blythe, 2005). The limestone in the Grand Wash trough is known as the Hualapai Limestone and has been correlated with similar limestone elsewhere in the eastern Lake Mead region (Longwell, 1928, 1936; Lucchitta, 1966). It has been interpreted as either marine (Blair, 1978; Blair and Armstrong, 1979) or nonmarine (Lucchitta, 1966; Faulds et al., 1997; Wallace, 1999). The rocks of the Grand Wash trough are bracketed between ca. 15 and 6 Ma. The older age is based on a 15.3 Ma 40Ar/39Ar date on sanidine from a rhyolite tuff near the base of the section on the west flank of Grapevine Mesa (Faulds et al., 2001b). Younger age constraints include (1) an 8.8 Ma basalt flow (Faulds et al., 2001b) intercalated with alluvial fan deposits shed from the
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Figure 6. Generalized geologic map for the Grand Wash trough, showing field trip stops. AF—Airport fault; AP—Airport Point; GC—Grapevine Canyon; GWC—Grand Wash Cliffs; ICF—Iceberg Canyon fault; LBF—Lost Basin Range fault; LP—Lookout Point; MF—Meadview fault; NGWF—northern Grand Wash fault; SCF—Sheep Canyon fault; SCVF—South Cove fault; SP—Sandy Point; WRF—Wheeler Ridge fault.
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A Grand Wash Trough
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Figure 7. Generalized stratigraphic columns of major basins near the eastern margin of the northern Colorado River extensional corridor, showing approximate tilts and relative unit thicknesses. (A) Grand Wash trough. QTc—Colorado River sediments; Tby—early Pliocene basalts (e.g., 4.4 Ma Sandy Point basalt; Th, 11–6 Ma Hualapai Limestone; Ts—Sandstone-siltstone facies; Tgy—gypsum; Tcg—conglomerate derived from Gold Butte block; Tbr—megabreccia of Proterozoic or Paleozoic rock; Ths—tuffaceous sedimentary rocks probably correlative with Horse Spring Formation (e.g., Beard, 1996) and containing a 15.3 Ma tuff (Faulds et al., 2001b); Mr—Redwall Limestone; Ds—Sultan Limestone; –Cm—Muav Limestone; –Cb—Bright Angel Formation; –Ct—Tapeats Sandstone; Xg—Paleoproterozoic gneiss. (B) Northern White Hills basin. See map (Fig. 10) for letter symbols. QTa—late Miocene-Quaternary basin-fill sediments; Tr—rounded gravel, sand, and silt; Th—Hualapai Limestone; Tc2—angular, poorly sorted sand and gravel, which has yielded dates from ca. 14.4–10.9 Ma (Blythe, 2005); Tb2—olivine basalt dated 8.3 Ma (Beard et al., 2007); Tb1—lower basaltic flows dated ca. 14.6 Ma (Duebendorfer and Sharp, 1998); Tc1—angular, poorly sorted sandy gravel (granular pattern) and lenses of megabreccia (blocky pattern); Tdt?—ash-flow tuff intercalated with megabreccia and poorly sorted angular conglomerate in Tc1, dated as 15.2 Ma (Duebendorfer and Sharp, 1998), may correlate with the tuff of Mt Davis (e.g., Faulds et al., 2002b); Tbs?—ash-flow tuff that may correlate with the tuff of Bridge Spring (e.g., Faulds et al., 2002b); Xg—Proterozoic gneiss. (C) Southern White Hills basin. Patterns and labels are the same as in Figure 11 except Tbr (megabreccia of Proterozoic rock). (D) Hualapai basin: QTa—Holocene-late Miocene shale, conglomerate, gypsum, and anhydrite; Tha—late Miocene (probably ca. 13–8 Ma) halite and lesser shale and anhydrite; Tcg—locally derived late Miocene fanglomerate; Tvsy—middle Miocene (ca. 16–13 Ma) volcanic and sedimentary rock; Tvso—early to middle Miocene (ca. 20–16 Ma) volcanic and sedimentary rock, possibly resting on a thin section of Cambrian strata; Xg—Proterozoic gneiss, granite, and diabase.
Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary Grand Wash Cliffs at Nevershine Mesa (Lucchitta et al., 1986) in the northern part of the trough; (2) a 7.43 ± 0.22 Ma 40Ar/39Ar maximum eruptive age on sanidine from an ash-fall tuff intercalated in the upper part of the Hualapai Limestone at Grapevine Mesa (Wallace et al., 2005); and (3) a 6.0 Ma tephra within the upper Hualapai Limestone in the Temple Bar area to the west of the Grand Wash trough (Spencer et al., 2001). Tilting within the rocks of the Grand Wash trough decreases up-section from ~30° in the lowermost ca. 15 Ma units to ~20 Ma
Ttb
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Tuff of Bridge Spring - 15.2 Ma
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Volcanics of Dixie Queen mine ~18.5-16.7 Ma dacite-andesite flows
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Figure 11. (A) Generalized geologic map of the southern White Hills and adjacent areas. (B) Cross section A–A′, shown slightly larger than map scale for ease of viewing but with no vertical exaggeration. The conglomerate of Rock Spring is a thick synextensional unit that thickens eastward in the eastern subbasin of the southern White Hills basin. Note that tilts progressively decrease upward between the upper part of the volcanics of Dolan Springs and basalt of Table Mountain Plateau. Greater tilting within the hanging wall indicates a listric geometry for the Cerbat Mountains and Cyclopic fault zones, which probably merge at depth beneath the eastern subbasin.
Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary Faulds et al., 2001b), which collectively mark the southern part of the South Virgin–White Hills detachment fault. As much as 3 km of Miocene volcanic and sedimentary strata accumulated on an erosion surface etched into Paleoproterozoic gneiss and granite, as well as Late Cretaceous peraluminous granite. In ascending order, the stratigraphy in the southern White Hills includes (Fig. 7C) (1) 18.5 to ca. 16 Ma trachydacite-trachyandesite lavas of the volcanics of Dixie Queen Mine (cf. Faulds et al., 1995); (2) basaltic andesite lavas of the volcanics of Dolan Springs; (3) a ca. 16 Ma bimodal sequence of intercalated rhyolite lavas, tuffs, and basaltic andesite flows, referred to as the volcanics of the southern White Hills; (4) the 15.2 Ma tuff of Bridge Spring (cf. Morikawa, 1994; Faulds et al., 1995, 2002b); (5) ca. 16–8.7 Ma synextensional fanglomerates, referred to as the conglomerate of Rock Spring; and (6) 8.7 Ma basalt of Table Mountain Plateau and 9.9 Ma basalt of Senator Mountain. The northerly striking Mountain Spring fault separates the southern White Hills into two distinct lithologic domains, or subbasins (Fig. 11). In the eastern subbasin, thick sections of fanglomerate (conglomerate of Rock Spring) and subordinate volcanic units accumulated in an eastward-thickening wedge bounded by the Cerbat Mountains and Cyclopic faults on the east. In contrast, volcanic rocks dominate the southern and central parts of the western subbasin, which developed in the mutual hanging walls of the Mountain Spring, Cerbat Mountains, and Cyclopic faults. As the volcanic section thins to the north in the northern part of the southern White Hills, however, the distinction between the eastern and western subbasins becomes less conspicuous. The large volcanic component in the western subbasin is more characteristic of half grabens within the bulk of the northern Colorado River extensional corridor (e.g., Anderson, 1971, 1978; Faulds, 1996; Faulds et al., 1995, 2001a, 2002b) and contrasts with the sediment-dominated basins along the eastern margin of the corridor and in the Lake Mead area (e.g., Lucchitta, 1966; Bohannon, 1984; Beard, 1996; Faulds et al., 1997, 2001c). The timing of extension within the southern White Hills is bracketed between ca. 16.7 and 8 Ma. Tilts within the southern White Hills progressively decrease up-section from ~75° in the volcanics of Dixie Queen Mine to ~5° in the basaltic lavas of Table Mountain Plateau (Fig. 12). Major east-west extension probably began ca. 16.7–16.2 Ma during deposition of the lowermost part of the conglomerate of Rock Spring, as evidenced by clasts of Proterozoic gneiss likely derived from surrounding footwall blocks. Although not fully exposed in any single fault block in the southern White Hills, concordant tilts in exposed parts of the lower Miocene section within individual fault blocks indicate little tilting and extension prior to ca. 16.7 Ma. Extension was clearly in full swing, however, during eruption of the ca. 16 Ma volcanics of the southern White Hills, as evidenced by angular unconformities with older units and appreciable tilt fanning (tilts decrease up-section from ~50° to 25°) (Figs. 11 and 12). Tilt fanning suggests that peak extension occurred between ca. 16.5 and 15 Ma. The final stages of extension are recorded by minor faulting and gentle tilting of the 8.7 Ma basalts. Although no evidence
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for Quaternary faulting was observed in the southern White Hills, it is possible that some of the faulting temporally correlates with that in the eastern Lake Mead area and may therefore be younger than ca. 6 Ma. The conglomerate of Rock Spring is a particularly important unit in the southern White Hills for elucidating the paleogeographic evolution of the region. It accumulated primarily in the eastern subbasin of the southern White Hills basin east of the Mountain Spring fault (Fig. 11). The conglomerate is bracketed between ca. 16.7 and 8.7 Ma by 40Ar/39Ar ages on underlying and overlying volcanic units. Clasts of Proterozoic gneiss and megacrystic granite, ranging up to ~10 m long, dominate the conglomerates. The matrix-supported texture, subangular clasts, and poor sorting suggest a fanglomerate origin for the conglomerate. The conglomerate forms an eastward-thickening wedge-shaped body in the southern White Hills and is as much as 2.6 km thick proximal to the Cerbat Mountains fault (Fig. 11). The conglomerate of Rock Spring probably correlates with similar conglomerate and rock avalanche deposits in the hanging wall of the northern White Hills basin, where Duebendorfer and Sharp (1998) documented thick middle Miocene rock-avalanche deposits of Proterozoic gneiss derived from the footwall of the detachment. Likely sources for the boulders of megacrystic Proterozoic granite in the conglomerate of Rock Spring include the southern footwall of the Cyclopic Mine fault and Garnet Mountain ~15 km to the east along the western margin of the Colorado Plateau (Figs. 2, 3, and 11). The relatively small body of megacrystic granite in the footwall of the Cyclopic fault, compared to the much larger body at Garnet Mountain, and the abundance of megacrystic granitic clasts within the entire 2.6-km-thick section of conglomerate of Rock Springs suggest that Garnet Mountain was a source for at least some of the detritus. Thus, the conglomerate of Rock Spring may record west-flowing drainages eroding headward into the footwall of the South Virgin–White Hills detachment fault and possibly into the western margin of the Colorado Plateau as early as ca. 16 Ma, essentially at the onset of major east-west extension. Hualapai Basin (Day 3) The Hualapai basin is a gently to moderately east-tilted half graben developed in the hanging wall of the southern Grand Wash fault (Figs. 2 and 13). Despite its proximity to the Colorado River and Grand Canyon, the Hualapai basin remains an internally drained, closed depression. Due to a lack of dissection by tributaries of the Colorado River, synextensional middle to late Miocene strata within the basin is obscured by more recent flatlying sediments (in contrast to the highly dissected Grand Wash trough and White Hills basin). Thus, both the stratigraphy and timing of extension cannot be directly inferred from exposures in the Hualapai basin. Nonetheless, drill-hole and seismic reflection data indicate that the Hualapai basin contains a thick (~3.9 km) growth-fault sequence of Miocene sedimentary and volcanic rocks, with tilts decreasing up section from ~25° to 0°. As inferred from analysis of core and seismic reflection profiles, the
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Figure 12. View to northeast of tilt fanning in southern White Hills basin. East tilts progressively decrease up-section from ~35° in 15.2 Ma tuff of Bridge Spring (Tbs) to ~5° in 8.7 Ma basalts of Table Mountain Plateau (Ttb). Trc—conglomerate of Rock Spring; Tv—volcanic rocks.
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Figure 13. Hualapai basin salt deposit. The Hualapai basin contains a 2.5-km-thick nonmarine salt deposit of probable late Miocene age, as evidenced by well and seismic reflection data. (A) 1:1 cross section constrained by a migrated seismic reflection profile (from Faulds et al., 1997). Patterns for units are the same as in Figure 7D. SGW, southern Grand Wash fault. (B) Core (9 cm in diameter) showing massive halite. (C) Core showing displacive halite crystals in reddishbrown claystone. This texture indicates deposition induced by groundwater discharge in a continental playa.
Cenozoic evolution of the abrupt Colorado Plateau–Basin and Range boundary stratigraphy of the Hualapai basin includes (in ascending order): (1) ~750 m of lower to middle Miocene volcanic and sedimentary rock possibly resting on Cambrian strata and/or Proterozoic gneiss, granite, and diabase; (2) ~335 m of middle Miocene volcanic and sedimentary rock; (3) fanglomerates along the margins that interfinger with evaporites in the central part of the basin; (4) up to 2500 m of middle to upper Miocene halite intercalated with minor shale (5%–10%) and anhydrite; and (5) ~600 m of late Miocene–Quaternary shale and lesser amounts of gypsum, anhydrite, and conglomerate (Faulds et al., 1997). The Miocene section in the Hualapai basin (Fig. 7D) is dominated by the 2.5-km-thick sequence of halite, one of the thickest known, nonmarine halite deposits in a continental rift (Faulds et al., 1997). The age of the salt deposit is roughly bracketed between ca. 13 and 8 Ma, because (a) it lies in the upper, more gently tilted part (18.56 Ma): Intermediate volcanism that erupted ignimbrites and possible lavas with coeval lahar and volcanogenic sediment deposition (dark gray layers and unit Tst in Fig. 6), all on Proterozoic granite basement. This was followed by a period of quiescence as rhythmically bedded sandstones (white stippled layers) that form the base of unit Tvs accumulated in a depositional basin. Time 2 (18.56 Ma): Rejuvenation of intermediate magmatism that resulted in the emplacement of trachydacite sills that intruded (and inflated) the earlier emplaced layered deposits. It is unclear if sill emplacement was associated with any eruptions. Time 3 (ca. 17.5–17.4 Ma): Eruption of ignimbrites and rhyolitic domes as well as rhyolite lava flows. Time 4 (ca. 17.4 Ma): Abrupt transition to intermediate magmatism marked by volcanic necks, dikes, and sills.
also include clasts of trachyandesite. In contrast, the block and ash flows consist of rhyolite clasts up to 20 cm in size, supported in a matrix dominated by ash and unflattened pumice. 40Ar/39Ar dating of rhyolite clasts from a block-and-ash-flow deposit and a rhyolite flow yield ages of 17.68 Ma (biotite) and 17.33 Ma (sanidine), respectively. SiO2 concentration in clasts from the breccia deposits range from 68 wt% to 73 wt%. Samples from the rhyolite lava flows (Trf) overlap in composition with the breccias but tend to be more silicic, ranging from 72 wt% to 82 wt% SiO2 (we infer that the most silicic samples have been altered, but some are probably true high-silica rhyolites). Numerous intrusions cross-cut units Tvs, Tbr, and Trf. These include a rhyolite dome (unit Tir), volcanic necks of trachydacite (unit Tid), and intermediate and felsic dikes and sills (unit Ti). The rhyolite dome is in the northwest corner of the map area, and the trachydacite necks (unit Ti) are to the southeast. The 40Ar/39Ar (biotite) age of one the necks is 17.45 Ma; thus, although it intrudes Tbr, it was essentially contemporaneous with the rhyolitic volcanism. For simplicity, the mafic to felsic dikes are grouped together. Rhyolite dikes are exposed only where they intrude Proterozoic granite (unit Xg), outside the fault block that is our focus. Dikes of
intermediate composition are aligned NW-SE, trending toward the trachydacite volcanic necks (Fig. 6), suggesting that they may all be part of a common feeder system. Where these dikes and related sills intrude into unit Tbr, they break up in peperitic fashion. This suggests that the breccias were wet, and probably young, when they were intruded by somewhat more mafic magmas. History of the Secret Pass Canyon Volcanic Center Based on the examination of field relations and petrographic, geochemical, and geochronologic data (Table 1) summarized above, we suggest the following history for the Secret Pass Canyon volcanic center (Fig. 8): 1. ≥18.55 Ma: Intermediate to felsic volcanism that erupted ignimbrites and lavas with coeval deposition of lahars and volcanogenic sediments (Tst). This was followed by a period of quiescence, as suggested by the rhythmically bedded brown sandstones at the base of unit Tvs. 2. 18.55 Ma: Emplacement of numerous trachydacite sills into Tst and the base of Tvs. The sills are interleaved and merge into a massive laccolith. No direct evidence that these trachydacite magmas reached the surface has been identified.
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Figure 9. Maps showing field trip stop locations. (A) Spirit Mountain Batholith (Day 1; taken from Google Earth). (B) Secret Pass Canyon volcanic center (Day 2). Dashed line with arrows in (B) refers to the walking route we will take.
Spirit Mountain batholith and Secret Pass Canyon volcanic center 3. From 18.55 to ca. 17.7 Ma: A possible period of quiescence within the volcanic center. This time may mark a period of magmatic differentiation from intermediate to felsic compositions. 4. Circa 17.7–17.4 Ma: Felsic eruptions: ignimbrites, airfall, dome formation, and collapse. 5. Circa 17.4 Ma: An abrupt transition to intermediate magmatism as marked by the volcanic necks as well as dikes and sills. FIELD TRIP STOPS (Fig. 9) Day 1: Spirit Mountain Batholith Mi Cumulative Description 0.0
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Leave hotel from the University of Nevada–Las Vegas; get on Tropicana, and start mileage. Merge onto 515/U.S. 93 (south). Follow the interstate out of Las Vegas and then merge onto U.S. 95 south. Turn left onto Christmas Tree Pass Road (east). Turn right onto dirt road heading up the hill (south). Stop 1: NO HAMMERS AT THIS STOP— this is sacred Native American land.
Stop 1: Spirit Mountain Batholith Roof We are now in the roof of the Spirit Mountain batholith, which is a ~25 km × 2 km zone of high-silica leucogranite. This zone comprises sheets of aplite, porphyry, and fine- to mediumgrained equigranular granite, with contacts that are sharp to barely perceptible. Most sheets were initially subhorizontal, but dikes (subvertical) are also common. We interpret these relations to indicate repeated emplacement of the leucogranites—some sheets intruding a hot, melt-bearing mush, and some intruding solid rock. Vesicles, or miarolitic cavities, are widespread and most common toward the west (top). Pegmatite pods and dikes dominated by coarse quartz and alkali feldspar are present locally. Typical leucogranites have ~40%–50% alkali feldspar (subhedral), 30%–40% quartz (anhedral in groundmass, but phenocrysts are subhedral), ~10% plagioclase (sub-euhedral), and ~1% biotite (euhedral). Porphyritic variants contain ~0.5 cm phenocrysts of quartz and alkali feldspar.
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Here you will see a large, rounded outcrop to your right (south); this is loosely known as the Catacombs. The road turns right (south), and you follow. Stop 2: NO HAMMERS AT THIS STOP; STILL SACRED LAND.
Stop 2: The Sacatone Wash–Quartz Monzonite Cumulate From the parking area, walk up the wash (west) for ~0.2 mi. The “wash” at this point turns into a steep-walled canyon. Continue on for ~0.1 mi to see good exposure of the quartz monzonite and giant mafic enclaves. The leucogranite from Stop 1 grades downward (eastward) through granite into magmatically foliated quartz monzonite, seen here, that is poorer in quartz and richer in biotite. The quartz monzonite is coarse-grained, with 40%–50% alkali feldspar (euhedral), 30%–35% plagioclase (euhedral), 10%–15% biotite (euhedral), 5%–15% quartz (interstitial, anhedral), and ~1%–2% hornblende (subhedral). Foliation, defined by aligned biotite, alkali feldspar, and plagioclase, is parallel to paleohorizontal and gradually becomes stronger downward. Dioritic enclaves are abundant, very large (up to 3m), pancake shaped, and oriented parallel to the rock’s ~N-S striking, W-dipping fabric (Fig. 10). Based on euhedral to subhedral feldspar crystal shapes and weakly to unstrained interstitial quartz, this fabric is interpreted to be dominantly magmatic and probably related to compaction of a crystal mush (cf. Bachmann and Bergantz, 2004). Mi Cumulative Description 3.0
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Get back on Christmas Tree Pass Road and drive south until the road dead ends into NV-163. Turn right (west) onto the highway. Follow the highway. Turn right (north) on a dirt road, which is across the highway
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Turn around and return to Christmas Tree Pass Road. Turn right (east). The dark colored outcrops to the left (north) are Precambrian country rock, consisting of megacrystic granite and orthogneiss.
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Figure 10. Dioritic enclaves within the leucogranite.
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from a large, sandy area used for off-road vehicle recreation. Road splits; stay right. Road crosses underneath a power line and across a power-line road. Continue straight. Continue straight on the road, keeping right if possible. This road can be fairly bad, so a 4WD truck is recommended. The road then turns sharply to the left (west). Don’t continue on, but park here. This is Bridge Canyon.
Stop 3: Bridge Canyon We are overlooking Bridge Canyon. Walk ~0.3 mi north, downhill into the canyon. Once in the canyon, go right (east), “downstream.” Here we see the granitic (sensu stricto) portion of the Spirit Mountain granite, some felsic dikes, and a few large pods of high silica granite segregate. The granite here is fairly homogeneous and is characterized by 30%–40% pink alkali feldspar (sub-euhedral), ~20%–35% plagioclase (euhedral laths), 15%–30% quartz (interstitial and anhedral to ~1 cm subhedral), 3%–8% biotite (euhedral), and ~1% sphene (euhedral). At this location, quartz forms prominent, round grains, but its abundance decreases markedly to the east. Locally, small (80% SiO2). Far to the southeast, you will see a sharp pyramidal peak that is surrounded by a blanket of scree (Fig. 15C). This peak is a volcanic neck of trachydacite (unit Ti) and is the collection location of geochron sample B on the geologic map (Fig. 6). Directly to the south are the rugged hills and dissected pediments of trachydacite that we interpret as a thick sill complex or laccolith (unit Td; Fig. 15D). Notice how monotonous this unit appears; we have not yet visited the highest points immediately south of the volcanic neck (unit Ti), but they appear to be continuous with the trachydacite complex. To the west across the Colorado River are Spirit Mountain and the Spirit Mountain Batholith, where we spent yesterday (Fig. 15E). At the base of the mountains is the east-dipping Newberry Detachment fault. You are currently standing on the upper plate of the detachment. Post-volcanic extension has tilted the Secret Pass Canyon volcanic center ~30°NE. At the bottom of the hill, turn left (north), retracing your route.
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Turn left (first left). We are driving along the east-dipping Arabian Mine Fault (measurements of the Proterozoic granite-Trachydacite contact indicate that the fault dips ~60° E); on the right (west) is Proterozoic granite (unit Xg), and on the left (east) is trachydacite (unit Td)]. Stay left on this road. Turn left (east) and follow the road up the wash (35°09′24.92′′ N, 114°27′40.77′′ W). During this drive, we pass through the trachydacite (unit Td), which again appears as a monotonous mass. The trachydacite continues to the south (right) until it is covered by Quaternary alluvium; reconnaissance to the south has not been able to constrain the bottom of this unit. Intersect with a north-south–trending road (35°9′22.08′′ N, 114°25′21.44′′ W). Our route to the south along this road, toward Oatman, Arizona, is hard to see—the road curves around the east side of a large outcrop of unit Td.
Stop 2: View of the Secret Pass Canyon Volcanic Center Stratigraphy To the north of this junction is a stair-stepped, flat-topped mountain that affords another view of the Secret Pass Canyon volcanic center stratigraphy (Fig. 16). The stratigraphy in this mountain is the same as seen within the crescent-shaped basin and ridge at Stop 1; the sun angle from about mid-afternoon to sunset highlights this stratigraphy magnificently. At the base of the mountain, the small rugged mounds mark the trachydacite (unit Td). The layered nonresistant section is the volcaniclastic sediments and reworked tuffs of unit Tvs. The breccias of unit
Figure 17. Photo of a mafic sill (unit Ti) that broke apart as it intruded into presumably wet breccias (unit Tbr). The sill broke apart into pillow-like features outlined with a white dashed line.
Tbr comprise the cliff-forming units within the middle, and a rhyolite flow of unit Trf marks the topmost point. Dark horizontal bands are exposed within the side of the mountain (in between units Tvs and Tbr and within unit Tbr) and represent more mafic sills. The topmost sill is broken apart within the breccias, suggesting intrusion into wet volcanogenic material, resulting in explosion features similar to a peperites (Fig. 17). If the sills did intrude into wet, unsolidified material, then it suggests that there was intermediate to mafic input into the magmatic system during or immediately subsequent to felsic volcanism. At this road junction, we are standing within the trachydacite (unit Td). Morphologically, the trachydacite has eroded into dissected hills that seem to have an almost a uniform slope of ~30°NE. Vertical jointing is also strongly apparent and, together with the uniform NE slopes, gives the impression that the trachydacite at this stop represents a suite of lava flows. However, we have found a distinct absence of obvious structures defining flows, such as basal brecciation, rubbly flow tops, and vesicles. Mi Cumulative Description 0.0
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Turn right (south; the road from this point on is extremely rough—4WD is necessary!). Along this road, we are completely within the trachydacite complex (unit Td). Pull off within the large wash and park (35°08′42.11′′ N, 114°25′25.38′′ W). As a point of reference, the peak on the east side of the road is labeled on the topographic map as having an elevation of 2540 ft.
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Stop 3: Road to Oatman, Arizona, at Peak 2540 We will be hiking within a wilderness area, so please proceed appropriately. At this stop, we will traverse up-section (east-northeast) through the trachydacite (unit Td) to what we interpret as an intrusive contact with unit Tvs; along the way, we will examine evidence that led to our interpretation of unit Td as an intrusive complex. We will then hike southeast within unit Tvs to the entrance to Secret Pass Canyon, where we will examine the spectacular cliffs comprised of avalanche breccias and block and ash flow deposits. We will then proceed southwest back down-section via Secret Pass Canyon wash. The following is a rough “walking log”; locations that we will spend significant time examining are labeled as stops 3a and 3b.
Figure 18. Fresh-cut hand sample of the trachydacite (unit Td) showing its coarse-grained nature. Phenocrysts comprise ~25%–35% of the sample and include plagioclase and biotite. Scale bar is 1 cm.
Trachydacite (Td)
Interfingered sediments (Tst)
Figure 19. Photo showing the general characteristics of the interfingered sediments that comprise unit Tst. The materials shown in this photo consist of a series of gray beds several centimeters thick composed of moderately to poorly sorted sand-sized grains of plagioclase, fresh-looking biotite, and lithics.
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(35°08′42.11′′ N, 114°25′25.38′′ W): We are beginning within the trachydacite section (unit Td) (Fig. 18). Plagioclase has weathered out in parts of the outcrop surface, giving the false appearance that the trachydacite contains vesicles; in fact, our inspection to date has yielded no vesicles throughout this section. As we walk up the wash, note on the lefthand side, close to the bottom of the outcrop, local lenses of rounded to subrounded trachydacite clasts that are on the order of centimeters in diameter. Compositionally, these clasts appear similar to their trachydacite host, which again appears to be massive and coarse-grained but are distinguishable from the host by slightly different weathering. These clasts represent either (1) material that was picked up during emplacement of the trachydacite, or (2) a brecciated flow base for the trachydacite. (35°08′50.37′′ N, 114°25′13.60′′ W): The wash splits here. Go to the right. (35°08′52.40′′ N, 114°25′5.95′′ W): The wash splits again. Go to the left. In this stretch, interfingered with trachydacite, we see volcanogenic sedimentary deposits. These sedimentary deposits define unit Tst in Figure 6 and outcrop in the walls of the wash as ~2–3-m-thick exposures that are capped by the trachydacite, which has slopes of ~30° NE. These units are gray in color, bedded on 1–10 cm scale. They contain moderately to poorly sorted grains of plagioclase, fresh-looking biotite, and lithic clasts up to tens of centimeters (Fig. 19); the fresh-looking biotite is easily distinguishable in hand samples and suggests that these deposits are locally derived. Several hundred meters farther up the wash (N35°08′52.26′′, 114°24′54.78′′W), we observe a poorly sorted breccia composed of subrounded to angular trachydacite and felsic (rhyolitic?) clasts up to >1 m in size (Fig. 20) that we interpret as a lahar deposit. Continuing up the wash, the volcanogenic deposits are again wellbedded, with fresh-looking biotite. It is worth noting that sedimentary deposits are not the only components of unit Tst. A welded tuff (possible rheomorphic ignimbrite?) is also exposed at 35°08′48.36′′ N, 114° 25′3.96′′ W.
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Figure 20. (A–C) Photos of the possible lahar deposits within unit Tst. The deposits shown in A–C are poorly sorted and composed of subrounded to angular trachydacite and felsic (rhyolitic?) clasts up to >1 m in size.
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Resistant ledge dipping ~30° NE
One trachydacite sill
Figure 21. Photo showing the typical slope characteristics of unit Td. As a whole, ridges that are capped by trachydacite appear to form laterally continuous layers that consistently dip ~30°NE.
A second trachydacite sill
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We are now back into the trachydacite. The trachydacite here is texturally, mineralogically, and compositionally similar to the trachydacite observed at the beginning of this traverse. If you climb on top of one of the walls of the wash, you begin to gain a sense that the trachydacites that cap all of the hills within this section have consistent ENE dip slopes. Further, it appears that, if not for erosion, the caps form laterally continuous layers traceable across much of unit Td (Fig. 21). These layers vary in their thickness from about meters to tens of meters. ~1.3 1.4 This is more of the trachydacite; however, blocks of bedded brown sedimentary material encased within the trachydacite become common in this stretch. These blocks are centimeters thick and vary from centimeters up to about a meter long with varying orientations within the trachydacite; the layers appear to have broken off along distinct bedding planes. Most appear as coherent blocks, but some appear to be warped, suggesting they have experienced ductile deformation. Stop 3a: Units Td-Tvs contact – 1.4 mi (35°9′5.76′′ N, 114°24′23.28′′ W): This is the contact between units Td and Tvs. At the base of the walls of the wash is trachydacite, which is ultimately overlain by rhythmically bedded brown sandstones that comprise the base of unit Tvs. The contact here, although roughly parallel to visible bedding, is locally irregular, with
fingers of the trachydacite intruding up into the sediments and along sediment bedding planes (Fig. 22). When observed along a bedding plane, the trachydacite appears as a brecciated sliver giving the impression that the trachydacite intruded wet sediment, causing it to break apart in a process akin to peperite formation (Fig. 22a). Where the trachydacite has intruded far enough along a bedding plane, blocks of the sediment have fallen—or been stoped —into the trachydacite. Again, these blocks have orientations that vary from parallel to the attitude of the coherent overlying brown sediments to blocks that are nearly perpendicular to this attitude. In many cases, the stoped blocks retain their primary sedimentary features, including bedding, cross bedding, and cut and fill structures. However, as noted in the previous section of this wash, some blocks have experienced ductile deformation. That blocks of overlying sedimentary material have been stoped into an underlying unit strongly indicates that the contact between units Td and Tvs at this locality is intrusive. The observation that the two units are roughly parallel to each other suggests that the intrusion was initially horizontal to subhorizontal— or sill-like—in nature. Based on the observations at this outcrop, together with the massive and coarse-grained nature of the trachydacite lower in the section and the absence of features defining flows, we infer an intrusive origin for unit Td. In this intrusive scenario, the volcanogenic sedimentary deposits observed lower in the section (Tst) are the remnants of the host rock into which the Td intruded. If this interpretation is correct, then the fact that unit Tst has an orientation consistent with the average attitudes at
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Figure 22. (A–B) Photos of what we interpret as an intrusive contact between units Td and Tvs. (A) Fingers of trachydacite that have intruded along individual bedding planes of the overlying brown sandstone have brecciated, giving them the appearance of peperites. (B) Blocks of the brown sandstone that have been stoped into the underlying trachydacite. Individual sandstone blocks have been rotated such that their orientations contrast with those of the overlying sediments. SS—sandstone; Td—trachydacite.
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the Secret Pass Canyon volcanic center suggest that these other intrusions were also sill-like. A test of this hypothesis would be dating the interfingered sedimentary deposits. If they yield ages >ca. 18.55 Ma (trachydacites are 18.55 Ma; Table 1), then the intrusive argument is strengthened. Mi Cumulative Description 1.4
~2.5
Climb out of the wash and hike to the southeast (paralleling the cliff face) toward the Secret Pass Canyon wash. The resistant cliffs to the east (left) are the breccias of unit Tbr. The less resistant layers below unit Tbr is the upper part of unit Tvs, which is composed of reworked airfall tuffs that are capped by an ignimbrite (Fig. 23). The reworked airfall tuffs consist of repeating sequences of red and white, moderately sorted, coarse- to fine-grained sandstones that are 3–10 cm thick and contain 1–2 cm diameter pieces of pumice (Fig. 23A), possibly reflecting Plinian-style activity here, which may have been a precursor to the eruption of the ignimbrite (Cas and Wright, 1987). The local ignimbrite unit is a 15–20-m-thick, pinkish, sanidine-bearing welded tuff with minor (3 m) quartz sandstone beds Slightly deeper cycles and more limestone than in adjacent sequences Carbonate cycles capped by thin (1.7-km-deep Late Devonian Alamo crater, southern Nevada [abs.]: Lunar and Planetary Institute Contribution No. 1234, abstract 1538, 2 p. Sandberg, C.A., Poole, F.G., and Morrow, J.R., 2006, Reinterpretation of Milk Spring and other oceanic deposits resulting from 382-Ma Alamo impact, southern Nevada, U.S.A., in Papers Presented to Conference on Impact
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Craters as Indicators for Planetary Environmental Evolution and Astrobiology, Östersund, Sweden, 2 p., http://www.geo.su.se/lockne2006/ AbstractVolume.htm. Sarg, J.F., 1988, Carbonate sequence stratigraphy, in Wilgus, C.K. et al., eds., Sea-level changes: An integrated approach: Society of Economic Paleontologists and Mineralogists Special Publication 42, p. 155–181. Schalla, R.A., and Benedetto, K.M., 1991, Early and Middle Devonian shelfslope transition, southern Mahogany Hills, Eureka County, Nevada: The Mountain Geologist, v. 28, no. 4, p. 13–23. Scholle, P.A., Bebout, D.G., and Moore, C.H., eds., 1983, Carbonate depositional environments: American Association of Petroleum Geologists Memoir 33, 708 p. Simo, J.A., Scott, R.W., and Masse, J.-P., eds., 1993, Cretaceous carbonate platforms: American Association of Petroleum Geologists Memoir 56, 479 p. Stevens, C.H., 1986, Evolution of the Ordovician through Middle Pennsylvanian carbonate shelf in east-central California: Geological Society of America Bulletin, v. 97, p. 11–25, doi: 10.1130/0016-7606(1986)972.0.CO;2. Stewart, J.H., 1972, Initial deposits of the Cordilleran geosyncline: Evidence of a late Precambrian (