Interior western United States (GSA Field Guide 6)

Interior western United States

Edited by Joel L. Pederson Utah State University Department of Geology 4505 Old Main Hill Logan, Utah 84322, USA and Carol M. Dehler Utah State University Department of Geology 4505 Old Main Hill Logan, Utah 84322, USA

Field Guide 6 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140 USA

2005

Copyright © 2005, 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 Interior western United States / edited by Joel L. Pederson and Carol Merritt Dehler. p. cm. — (Field guide ; 6) Includes bibliographical references. ISBN 0-8137-0006-X (pbk.) 1. Geology--West (U.S.)--Guidebooks. 2. West (U.S.)--Guidebooks. I. Pederson, Joel L., 1968II. Dehler, Carol Merritt, 1964- III. Field guide (Geological Society of America) ; 6. QE79.I567 2005 557.88-dc22 2005052188 Cover: View looking from Muley Point across part of the Goosenecks of the San Juan River (entrenched meanders) in southeastern Utah. The escarpment in the left foreground is Permian Cedar Mesa Sandstone and the upper canyon walls of the goosenecks are the Pennsylvanian Honaker Trail Formation. Photograph by Tim Demko.

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Contents

1. Neoproterozoic Uinta Mountain Group of northeastern Utah: Pre-Sturtian geographic, tectonic, and biologic evolution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 Carol M. Dehler, Douglas A. Sprinkel, and Susannah M. Porter 2. Basaltic volcanism of the central and western Snake River Plain: A guide to field relations between Twin Falls and Mountain Home, Idaho. . . . . . . . . . . . . . . . . 27 John Shervais, John D. Kauffman, Virginia S. Gillerman, Kurt L. Othberg, Scott K. Vetter, V. Ruth Hobson, Meghan Zarnetske, Matthew F. Cooke, Scott H. Matthews, Barry B. Hanan 3. From cirques to canyon cutting: New Quaternary research in the Uinta Mountains. . . . . . . . . 53 Jeffrey S. Munroe, Benjamin J.C. Laabs, Joel L. Pederson, and Eric C. Carson 4. Geomorphology and rates of landscape change in the Fremont River drainage, northwestern Colorado Plateau . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79 David W. Marchetti, John C. Dohrenwend, and Thure E. Cerling 5. Late Cretaceous stratigraphy, depositional environments, and macrovertebrate paleontology of the Kaiparowits Plateau, Grand Staircase–Escalante National Monument, Utah. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101 Alan L. Titus, John D. Powell, Eric M. Roberts, Scott D. Sampson, Stonnie L. Pollock, James I. Kirkland, and L. Barry Albright 6. Transect across the northern Walker Lane, northwest Nevada and northeast California: An incipient transform fault along the Pacific–North American plate boundary . . . . . . . . . . . 129 James E. Faulds, Christopher D. Henry, Nicholas H. Hinz, Peter S. Drakos, and Benjamin Delwiche 7. Brittle deformation, fluid flow, and diagenesis in sandstone at Valley of Fire State Park, Nevada . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 151 Peter Eichhubl and Eric Flodin 8. Evolution of a late Cenozoic supradetachment basin above a flat-on-flat detachment with a folded lateral ramp, SE Idaho . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 Alexander N. Steely, Susanne U. Janecke, Sean P. Long, Stephanie M. Carney, Robert Q. Oaks, Jr., Victoria E. Langenheim, and Paul K. Link 9. Utah’s state rock and the Emery coalfield: Geology, mining history, and natural burning coal beds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 199 Glenn B. Stracher, David E. Tabet, Paul B. Anderson, and J. Dénis N. Pone

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10. Latest Pleistocene–early Holocene human occupation and paleoenvironmental change in the Bonneville Basin, Utah–Nevada. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 211 David Rhode, Ted Goebel, Kelly E. Graf, Bryan S. Hockett, Kevin T. Jones, David B. Madsen, Charles G. Oviatt, and Dave N. Schmitt 11. Neotectonics and paleoseismology of the Wasatch Fault, Utah . . . . . . . . . . . . . . . . . . . . . . . . . 231 Ronald L. Bruhn, Christopher B. DuRoss, Ronald A. Harris, and William R. Lund 12. Pocatello Formation and overlying strata, southeastern Idaho: Snowball Earth diamictites, cap carbonates, and Neoproterozoic isotopic profiles . . . . . . . . . . . . . . . . . . . . . . . 251 Paul Karl Link, Frank A. Corsetti, and Nathaniel J. Lorentz 13. Anatomy of reservoir-scale normal faults in central Utah: Stratigraphic controls and implications for fault zone evolution and fluid flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 261 Peter Vrolijk, Rod Myers, Michael L. Sweet, Zoe K. Shipton, Ben Dockrill, James P. Evans, Jason Heath, and Anthony P. Williams 14. Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths of the Henry Mountains, southern Utah . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 283 Sven Morgan, Eric Horsman, Basil Tikoff, Michel de Saint-Blanquat, and Guillaume Habert 15. Folds, fabrics, and kinematic criteria in rheomorphic ignimbrites of the Snake River Plain, Idaho: Insights into emplacement and flow . . . . . . . . . . . . . . . . . . . . . 311 Graham D.M. Andrews and Michael J. Branney 16. Mesozoic lakes of the Colorado Plateau . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 329 Timothy M. Demko, Kathleen Nicoll, Joseph J. Beer, Stephen T. Hasiotis, and Lisa E. Park 17. Birth of the lower Colorado River—Stratigraphic and geomorphic evidence for its inception near the conjunction of Nevada, Arizona, and California. . . . . . . . . . . . . . . . . . . 357 P. Kyle House, Phillip A. Pearthree, Keith A. Howard, John W. Bell, Michael E. Perkins, James E. Faulds, and Amy L. Brock 18. Development of Miocene faults and basins in the Lake Mead region: A tribute to Ernie Anderson and a review of new research on basins . . . . . . . . . . . . . . . . . . . . 389 Melissa Lamb, Paul J. Umhoefer, Ernie Anderson, L. Sue Beard, Thomas Hickson, and K. Luke Martin 19. Don R. Currey Memorial Field Trip to the shores of Pleistocene Lake Bonneville . . . . . . . . . 419 Holly S. Godsey, Genevieve Atwood, Elliott Lips, David M. Miller, Mark Milligan, and Charles G. Oviatt 20. Paleoseismology and geomorphology of the Hurricane Fault and Escarpment . . . . . . . . . . . . 449 Lee Amoroso and Jason Raucci 21. Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior, Book Cliffs, Utah . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 479 Simon A.J. Pattison 22. Geologic hazards of the Wasatch Front, Utah . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 505 Barry J. Solomon, Francis X. Ashland, Richard E. Giraud, Michael D. Hylland, Bill D. Black, Richard L. Ford, Michael W. Hernandez, and David H. Hart

Geological Society of America Field Guide 6 2005

Neoproterozoic Uinta Mountain Group of northeastern Utah: Pre-Sturtian geographic, tectonic, and biologic evolution Carol M. Dehler Department of Geology, Utah State University, 4505 Old Main Hill, Logan, Utah 84322, USA Douglas A. Sprinkel Utah Geological Survey, P.O. Box 14610, Salt Lake City, Utah 84114, USA Susannah M. Porter Department of Earth Science, University of California–Santa Barbara, Santa Barbara, California 93106, USA

ABSTRACT The Neoproterozoic Uinta Mountain Group is undergoing a new phase of stratigraphic and paleontologic research toward understanding the paleoenvironments, paleoecology, correlation across the range and the region, paleogeography, basin type, and tectonic setting. Mapping, measured sections, sedimentology, paleontology, U-Pb geochronology, and C-isotope geochemistry have resulted in the further characterization and genetic understanding of the western and eastern Uinta Mountain Group. The Red Pine Shale in the western Uinta Mountain Group and the undivided clastic strata in the eastern Uinta Mountain Group have been a focus of this research, as they are relatively unstudied. Reevaluation of the other units is also underway. The Red Pine Shale is a thick, organic-rich, fossiliferous unit that represents a restricted environment in a marine deltaic setting. The units below the Red Pine Shale are dominantly sandstone and orthoquartzite, and represent a fluviomarine setting. In the eastern Uinta Mountain Group, the undivided clastic strata are subdivided into three informal units due to a mappable 50–70-m-thick shale interval. These strata represent a braided fluvial system with flow to the southwest interrupted by a transgressing shoreline. Correlation between the eastern and western Uinta Mountain Group strata is not complete, yet distinctive shale units in the west and east may be correlative, and one of the latter has been dated (≤770 Ma). Regional correlation with the 770–742 Ma Chuar Group suggests the Red Pine Shale may also be ca. 740 Ma, and correlation with the undated Big Cottonwood Formation and the Pahrump Group are also likely based upon C-isotope, fossil, and provenance similarities. This field trip will examine these strata and consider the hypothesis of a ca. 770–740 Ma regional seaway, fed by large braided rivers, flooding intracratonic rift basins and recording the first of three phases of rifting prior to the development of the Cordilleran miogeocline. Keywords: Neoproterozoic, Uinta Mountain Group, intracratonic rift, vase-shaped microfossil, Bavlinella faveolata, Leiosphaeridia sp. Dehler, C.M., Sprinkel, D.A., and Porter, S.M., 2005, Neoproterozoic Uinta Mountain Group of northeastern Utah: Pre-Sturtian geographic, tectonic, and biologic evolution, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 1–25, doi: 10.1130/ 2005.fld006(01). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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INTRODUCTION This field trip guide is a review and an update of the existing data sets regarding Uinta Mountain Group geology and reports the latest ideas about depositional environments, correlation, paleogeography, biologic evolution, and tectonic setting in northeastern Utah during Neoproterozoic time. The focus of the field trip will be on the previously understudied units (Red Pine Shale and undivided eastern clastic strata), as well as a reevaluation of previous interpretations about the other better-studied units, and how all of these units relate to one another in time and space. The Uinta Mountain Group is interesting for many reasons: (1) it is one of few exposed strata in the region for understanding the early tectonic evolution of the Late Neoproterozoic western Laurentian margin; (2) it likely records the inception of climate change leading into the low-latitude glaciations of the Sturtian episode; and (3) it contains a wealth of microfossils throughout the succession that can inform us of pre-Sturtian biologic evolution and how it may relate to (1) and (2) above. Lastly, (4) the Uinta Mountain Group is a “sleeping giant” in terms of being explored for its information on Precambrian geology. It has been under the “curse of the Proterozoic sandstones” (Link et al., 1993) for too long, and ongoing and future research will hopefully lift the curse. A general overview of the Uinta Mountain Group is provided first, followed by a two-day road log that takes the reader clockwise around the Uinta Mountains. On Day 1, the western and central strata of the north flank will be visited, and on Day 2 the easternmost strata and the strata of the south flank will be viewed and discussed.

tion indices (TAI) indicate 4.5–7 km of strata (Hansen, 1965; Stone, 1993; Sprinkel et al., 2002). It is uncertain how the eastern and western Uinta Mountain Group strata correlate; however, similar petrographic patterns are evident across the range and throughout the group. Geochemical and provenance studies show arkosic sandstone and shale in the north part of the range were derived from the Wyoming craton to the north, and quartz arenite in the southern part of the range was derived, in part, from a Paleoproterozoic source to the east (e.g., Wallace, 1972; Sanderson, 1978; Sanderson, 1984; Ball and Farmer, 1998; Condie et al., 2001). The age of the Uinta Mountain Group is likely entirely Neoproterozoic. It unconformably overlies the metamorphic quartzitic and schistose units of the Red Creek Quartzite and Owiyukuts Complex (ca. 1.7 to ca. 2.7 Ga; Hansen, 1965; Sears et al., 1982) and is unconformably overlain by lower Paleozoic strata. A 770 Ma detrital zircon population from the middle eastern Uinta Mountain Group (Fanning and Dehler, 2005) indicates that the majority of the Uinta Mountain Group is younger than 770 Ma. The uppermost unit in the Uinta Mountain Group, the Red Pine Shale, yielded a microfossil assemblage and C-isotope variability similar to that of the 742 Ma upper Chuar Group in Arizona, therefore putting a possible upper age limit on the Uinta Mountain Group (Vidal and Ford, 1985; Karlstrom et al., 2000; Porter and Knoll, 2000; Dehler, 2001; Dehler et al., 2006). Paleomagnetic data from the Uinta Mountain Group indicate deposition in equatorial latitudes, and the Uinta Mountain Group paleopole sits right on the Chuar Group apparent polar wander path, also suggesting a similar age (Weil et al., 2005). Western Uinta Mountain Group Stratigraphy

UINTA MOUNTAIN GROUP STRATIGRAPHY The Neoproterozoic Uinta Mountain Group is a 4–7-kmthick siliciclastic succession that is exposed only in the Uinta Mountains and makes up the core of the Uinta Mountain anticline (Fig. 1). The strata exposed in the western Uinta Mountains are characteristically different than those in the eastern Uinta Mountains, perhaps due to structural subbasins imparting control on depositional style. Hansen (1965) indentified two structural domes within the overall Uinta anticline, one in the western and one in the eastern part of the range, and these roughly correspond to the changes in stratal character. In the western Uinta Mountains, the Uinta Mountain Group comprises >4 km of sandstone and sedimentary quartzite, with lesser shale and rare conglomerate (Wallace, 1972) (Figs. 1 and 2). These strata show much lateral and vertical variability and have undergone many subdivisions (see Sanderson, 1984). In the eastern part of the range, the Uinta Mountain Group is dominantly sandstone with lesser shale and a distinctive basal conglomerate and breccia (Jesse Ewing Canyon Formation; Sanderson and Wiley, 1986). The base of the Uinta Mountain Group is exposed only in the eastern Uinta range, and calculated thicknesses from air photos, seismic profiles, and thermal altera-

The western Uinta Mountain Group has been subdivided several different ways (Williams 1953; Wallace and Crittenden, 1969; Wallace, 1972; Sanderson, 1984). The nomenclature used here is mainly after Wallace (1972), in combination with what has been considered mappable on a 1:125,000 scale by Bryant (1992). These units include the lowermost formation of Moosehorn Lake (including the basal undivided Uinta Mountain Group), the formation of Red Castle, the formation of Dead Horse Pass, the Mount Watson Formation, the formation of Hades Pass, and the Red Pine Shale (Figs. 2 and 3). Only the Mount Watson Formation has been formally named following the Stratigraphic Code (Sanderson, 1984). The Red Pine Shale was formalized by Williams (1953) prior to adoption of the Stratigraphic Code (North American Stratigraphic Commission on Nomenclature, 1983). Other informal units proposed by Wallace (1972) may be mappable at larger scales, but will not be featured in this paper. Basal Undivided Uinta Mountain Group and Formation of Moosehorn Lake The lowermost 60 m of exposed Uinta Mountain Group in the western range is an undivided interval of white quartz arenite. It is likely a lateral equivalent of the formation of Red Castle, and

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TIntic Quartzite Mutual Formation Mineral Fork Tillite Big Cottonwood Formation Little Willow Series Farmington Canyon Complex

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Red Pine Shale

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Red Pine Shale Uinta Mountain Group, Undivided Jesse Ewing Canyon Formation Red Creek Quartzite And Owiyukuts Complex

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Lodore Formation Cl

UTAH

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Figure 1. Geologic map of the Uinta Mountains and the adjacent Wasatch Range with an emphasis on Precambrian geology, showing field trip stops for Day 1 and Day 2.

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Neoproterozoic Uinta Mountain Group of northeastern Utah 3

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fm. of Outlaw Trail

fm. of 1 Diamond Breaks south of Browns Park. 4-5 km of partially subdivided UMG, base not exposed

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north of Browns Park, ~7 km of undivided UMG overlies basal Jesse Ewing Canyon Fm. sh s

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Uinta Mountain Group Eastern Uinta Mountains

Figure 2. Stratigraphic columns for the western and eastern Uinta Mountain Group (UMG). Data for the western Uinta Mountain Group stratigraphic column is from Wallace (1972), Sanderson (1978, 1984), and Dehler et al. (2006). Data for the eastern Uinta Mountain Group is from Hansen (1965), Sanderson and Wiley (1986), Hansen and Rowley (1991), De Grey (2005), Nagy and Porter (2005), and Dehler et al. (2006). Note that there is a break in section in the eastern Uinta Mountain Group column. The Jesse Ewing Canyon Formation is overlain by ~7 km of undivided Uinta Mountain Group strata and this succession is exposed solely on the north side of the Browns Park graben. On the south side of the Browns Park graben, the divided eastern Uinta Mountain Group is exposed, although neither the basal contact nor the Jesse Ewing Canyon Formation are exposed there. Future work will determine how these strata correlate across the graben and how the eastern and western Uinta Mountain Group strata correlate across the range.

is only exposed in the southern and central parts of the western Uinta range (Fig. 1) (Wallace 1972). The base of this unit is not exposed, and it grades upward into the formation of Moosehorn Lake. In the mapping of Bryant (1992) and in this paper, this unit is included in the basal part of the formation of Moosehorn Lake (Figs. 2 and 3). The formation of Moosehorn Lake (~150–300 m thick) is a dark olive-green to yellow green shale with thin lenticular to tabular interbeds of pebbly arkosic arenite (Fig. 4). Common sedimentary features include ripplemarks, mudcracks, and softsediment deformation (Wallace and Crittenden, 1969; Wallace, 1972). It is exposed in the Bald Mountain area of the higher western Uinta Mountains (Fig. 2). It is overlain by the Mount Watson

Formation or formation of Dead Horse Pass, and northward it grades into and is overlain by the lower part of the formation of Red Castle (Fig. 4) (Wallace and Crittenden, 1969; Wallace, 1972). The formation of Moosehorn Lake was interpreted to represent a spectrum of marine environments (see Day 1, Stop 2) (Wallace, 1972). The undivided underlying quartz arenite was interpreted to represent a braided stream system that is related to the stream system in the formation of Red Castle. Formation of Red Castle The formation of Red Castle (>730 m thick) comprises dominantly arkosic arenite with subordinate subarkosic and quartz arenite. Common sedimentary features include trough and

Neoproterozoic Uinta Mountain Group of northeastern Utah

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Deseret Formation

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formation of Outlaw Trail formation of Diamond Breaks

formation of Moosehorn Lake

? E EARLY PROTEROZOIC AND ARCHEAN

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WESTERN WEST-CENTRAL EAST-CENTRAL SOUTH-EASTERN NORTH-EASTERN UINTA MOUNTAINS UINTA MOUNTAINS UINTA MOUNTAINS UINTA MOUNTAINS UINTA MOUNTAINS

UMG undivided

Jesse Ewing Canyon Formation

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Red Creek Quartzite and Owiyukuts Complex?

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Red Creek Quartzite and Owiyukuts Complex?

Red Creek Quartzite and Owiyukuts Complex

Figure 3. Correlation chart for the Uinta Mountain Group. Stratigraphic nomenclature is from Wallace (1972), Sanderson (1978, 1984), De Grey (2005) and De Grey and Dehler (2005). The distribution of formations is based on mapping by Bryant (1992), Sprinkel (2002), and De Grey (2005). The 770 Ma detrital zircon age is from Fanning and Dehler (2005) and the ca. 740 age of the top of the Red Pine Shale is based on correlation with the 742 Ma Chuar Group (Dehler et al., 2006).

Figure 4. The east face of Bald Mountain along the Mirror Lake Scenic Byway. Bald Mountain is comprised of the Mountain Watson Formation, which is thick- to medium-bedded quartz arenite and sedimentary quartzite; note the massive-weathering cliffs near the base and top of the mountain with an intervening slope-forming unit. The underlying formation of Moosehorn Lake forms the road cut above the vehicle (near center of photo). The bed top in the foreground is a medium-bedded sandstone interbed within the formation of Moosehorn Lake. The contact between the formation of Moosehorn Lake and the Mountain Watson Formation is at the base of the lower massiveweathering cliff near the bottom of Bald Mountain. Photo by Bart Kowallis.

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planar-tabular cross-bedding, mudcracks, ripplemarks, and mud chips (Wallace, 1972). It is exposed in the higher Uinta Mountains from the area of Red Castle Lake west to the area of Dead Horse and Amethyst lakes (Fig. 1). This unit is conformably overlain by the formation of Hades Pass. The formation of Red Castle is interpreted to represent a braided stream system. Paleocurrent analysis indicates flow from the north, northeast, and east (Wallace and Crittenden, 1969; Wallace, 1972). Mount Watson Formation The Mount Watson Formation (~550–1000 m thick) comprises dominantly gray to white quartz arenite and lesser subarkosic arenite, with subordinate lenticular gray-green shale and reddish arkosic arenite interbeds (see Day 1, Stop 2). It is exposed in the Mount Watson area of the high western Uinta Mountains (Bryant, 1992; Fig. 1). It is laterally gradational with the formation of Dead Horse Pass and is overlain by the formation of Hades Pass, both in gradational contact (Figs. 4 and 5) (Wallace, 1972; Sanderson, 1978, 1984). The Mount Watson Formation was interpreted to represent a mixture of fluvial, coastal, and marine environments by Wallace (1972) and to represent a braided sheetwash plain by Sanderson (1978, 1984). Formation of Dead Horse Pass This formation (900 m thick) is dominated by quartz arenite and orthoquartzite, with subordinate shale and siltstone. Common sedimentary features include “giant planar crossbed sets,” planar-tabular and trough crossbeds, thin beds of alternating sandstone and shale, ripplemarks, mudchips, and soft-sediment deformation. Shale intervals are locally interbedded with thin beds of sandstone. Shale beds exhibit mudcracks, interference ripples, terraced current ripplemarks, and mudcracks (Wallace, 1972; Sanderson, 1978; Sanderson, 1984). Shale intervals are tens to hundreds of meters thick; the thickest shale interval (~200 m) is informally named the Gilbert Peak Shale member by Wallace (1972). This unit is exposed in the central higher Uinta Mountains in the Dead Horse Pass area at the head of Rock Creek (Fig. 1) (Wallace, 1972). Using the mapping divisions of Bryant (1992), this latter formation includes the Gilbert Peak Shale member and the Mount Agassiz formation of Wallace (1972). Where this unit becomes shale- and siltstone-poor and rich in subarkose arenite to the west, it is called the Mount Watson Formation (Fig. 3). To the north, this unit grades into a poorly sorted arkosic arenite and is called the formation of Red Castle (Wallace, 1972). Paleoenvironmental interpretations are similar to those of the Mount Watson Formation. Wallace (1972) interpreted this unit to represent mostly coastal and marine environments, and Sanderson (1978, 1984) interpreted this unit as a braided sheetwash plain. Formation of Hades Pass The formation of Hades Pass (1825–3600 m thick) comprises red to purple quartz arenite, subarkosic arenite, and arkosic arenite (Fig. 5). It has subordinate shale intervals that are grayish red, grayish olive green, or yellow and are ≤50 m thick.

Sedimentary structures in the sandstone units include medium to thick beds of planar-tabular and trough crossbedding (some indicating reverse current flow) and soft-sediment deformation (Wallace and Crittenden, 1969; Wallace, 1972). It is recognized in the central and western parts of the higher Uinta Mountains (Fig. 1). The upper contact with the overlying Red Pine Shale is gradational (Wallace and Crittenden, 1969; Wallace, 1972). This unit is interpreted to represent a probable fluvial origin. Paleocurrent data indicate a dominantly east-west flow direction, yet the dominant sandstone type is arkosic arenite. Interestingly, in the westernmost exposure of this unit, paleocurrent directions show dominantly northwest paleocurrent flow (Wallace and Crittenden, 1969; Wallace, 1972). Red Pine Shale The Red Pine Shale (Williams, 1953) comprises organicrich gray shale, siltstone, and subordinate sandstone (quartz arenite to arkosic arenite) (see Day 1, Stops 1 and 3; Day 2, Stops 5 and 6) (Fig. 6). It is overlain unconformably by the locally laterally discontinuous Cambrian Tintic Quartzite or the overlying Mississippian Madison Formation (Figs. 1 and 2). It ranges in measured thickness from ~300 m to >1200 m on the south flank and is between ~500 m and ~1825 m thick on the north flank (Williams, 1953; Wallace, 1972; Bryant, 1992; Dehler et al., 2006). The facies characteristics and associations indicate offshore deposition near or below fair-weather wavebase, in a deltaic system (Dehler et al., 2006). Eastern Uinta Mountain Group Stratigraphy Stratigraphic research on the eastern Uinta Mountain Group has resulted in the subdivision of the basal ~225 m into the Jesse Ewing Canyon Formation (Sanderson and Wiley, 1986) and the new division of the majority of correlative overlying strata into the formations of Diamond Breaks, Outlaw Trail, and Crouse Canyon and undivided Uinta Mountain Group (Figs. 2 and 3) (e.g., 1:24,000 scale mapping; De Grey, 2005; Dehler et al., 2006). Constraints on the thickness of the eastern Uinta Mountain Group are poorly known, yet seismic data interpretation, thermal maturation data, and mapping suggest that the Uinta Mountain Group is ~4.5–7 km thick in the northernmost area of exposure (juxtaposed to the Uinta Fault zone) and between 4 and 5 km southward (Fig. 2) (Hansen, 1965; Stone, 1993; De Grey, 2005; Sprinkel and Waanders, 2005). Jesse Ewing Canyon Formation The Jesse Ewing Canyon Formation (~225 m thick) comprises lithic clast-supported conglomerate and breccia, lithic and quartz arenite, and red to black shale (Fig. 7; see Day 2, Stop 2) (Sanderson and Wiley, 1986). Abrupt north-to-south facies changes show coarse conglomerate and breccia beds thinning southward into thick intervals of gray to red to green shale and subordinate interbeds of sandstone (Sanderson and Wiley, 1986; Dehler et al., 2006). Paleocurrent data measured

Neoproterozoic Uinta Mountain Group of northeastern Utah

Figure 5. View northeast of Hayden Peak (12,479 ft). The formation of Hades Pass (upper, darker unit) caps the peak with most of the mountain comprising the underlying lighter-colored Mount Watson Formation. Hayden Peak is faulted on the north and south. The rocks that form Hayden Peak dip northward. The axis of the Uinta arch is mapped by Bryant (1992) approximately through the Hayden Peak overlook stop.

from different types of crossbedding in interbedded sandstone units indicate a southwesterly flow direction (Sanderson and Wiley, 1986). The Jesse Ewing Canyon Formation unconformably overlies the Paleoproterozoic-Archean(?) Red Creek Quartzite and is overlain in gradational contact by undivided Uinta Mountain Group (Sanderson and Wiley, 1986) (Fig. 3). It is exposed in a ~56 km2 area in the Jesse Ewing Canyon area, south of Clay Basin and north of Browns Park, very near and west of the Utah-Colorado border (Fig. 1). This unit is truncated to the south by the Tertiary Mountain Home fault, part of the Browns Park graben, and is not found southward, nor is the base of the Uinta Mountain Group exposed elsewhere except in this area north of the graben (Fig. 1). The Jesse Ewing Canyon formation laterally pinches out to the west and east, and, in these areas, undivided Uinta Mountain Group sandstone rests directly on the Red Creek Quartzite (Hansen, 1965; Sprinkel, 2002). Sanderson and Wiley (1986) suggested that this unit represents alluvial fan and related deposits. Dehler et al. (2006) suggest that parts of this unit indicate subaqueous-mass-flow and/or fan delta deposition along a wave-affected shoreline. Uinta Mountain Group Undivided above Jesse Ewing Canyon Formation Approximately 7 km of sandstone and interbedded shale conformably overlie the Jesse Ewing Canyon Formation (Hansen, 1965) and have yet to be subdivided or measured and described in detail. These strata, and the underlying Jesse Ewing Canyon Formation, are stratigraphically isolated on the north side of the Browns Park graben from newly subdivided Uinta Mountain Group strata on the south side of the graben

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Figure 6. View of a north-facing outcrop of the Red Pine Shale from the Castle Rocks overlook. Hades Creek is at the base of the outcrop. Note sandstone interbeds toward top of exposure. This unit is interpreted to represent a marine-deltaic system.

Figure 7. Photo of Jesse Ewing Canyon Formation showing the lateral facies change from predominantly conglomerate and breccia to the north to dominantly shale to the south. Although previously interpreted as an alluvial fan deposit, a significant amount of the Jesse Ewing Canyon Formation was deposited in a subaqueous environment, possibly marine.

(Figs. 1, 2, and 3). This undivided Uinta Mountain Group unit consists of pebbly sandstone interbedded with red shale intervals. The sandstone intervals are typically tens of meters thick, and the shale intervals are meters thick. Common sedimentary features are trough crossbeds and soft-sediment deformation. Hansen (1965) interpreted these strata to represent shallow- to marginal-marine and subaerial environments in a rapidly subsiding trough.

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Figure 8. View looking south at the Diamond Breaks on the south side of Browns Park. The white lines indicate contacts between the three informal units of the eastern Uinta Mountain Group proposed by De Grey (2005). Abbreviations denote Neoproterozoic formations of Diamond Breaks (Zud), Outlaw Trail (Zuo), and Crouse Canyon (Zuc).

Formation of Diamond Breaks The formation of Diamond Breaks (500–1000 m thick) is the lowermost subdivided informal unit in the undivided eastern Uinta Mountain Group, and it is not yet certain how this unit correlates with the undivided Uinta Mountain Group and Jesse Ewing Canyon Formation to the north (Figs. 1 and 2). This formation comprises dominantly quartz arenite, with subordinate arkosic and subarkosic arenite with subordinate thin intervals of red shale (see Day 2, Stop 3) (De Grey, 2005). It is sharply, but conformably, overlain by the formation of Outlaw Trail, and the base is not exposed (Fig. 8). The formation of Diamond Breaks contains facies associations that represent various depositional environments within a braided river system. Formation of Outlaw Trail The formation of Outlaw Trail (50–70+ m thick) comprises green to gray to red shale, interbedded with thin to thick arkosic sandstone beds (Fig. 9) (see Day 2, Stop 3). It is exposed along the north face of the Diamond Breaks and can be traced for tens of kilometers laterally (Fig. 8) (Dehler et al., 2006). The formation of Outlaw Trail has been interpreted to be the low energy, interdistributary area of a proximal to medial delta plain environment, such as a bay, lagoon, swamp, or lacustrine environment (Dehler et al., 2006). Formation of Crouse Canyon The formation of Crouse Canyon sharply overlies the formation of Outlaw Trail (Fig. 8). The thickness of the formation of Crouse Canyon in this area reaches up to 1170 m and is estimated to be ~3200 m if extended to the top of the Uinta Mountain Group (Fig. 2) (De Grey and Dehler, 2005). This formation is similar to the formation of Diamond Breaks (see Day 2, Stops 3 and 4). Depositional environments are similar to those in the formation of Diamond Breaks. Undivided Uinta Mountain Group above formation of Crouse Canyon. The upper 2030 m of the eastern Uinta Mountain

Group have not been described in detail, and are here included in the formation of Crouse Canyon. Preliminary observation and geologic mapping indicate that these strata are very similar to the underlying formation of Crouse Canyon. This unit also very likely represents a braided stream environment (De Grey, 2005). Paleontology of the Uinta Mountain Group Overview The Uinta Mountain Group was deposited during a transitional time in the early evolution of the biosphere. After a long interval characterized by limited diversity and low abundance, eukaryotes were diversifying and expanding into prokaryotedominated environments. Although animals had not yet originated, protistan clades including red algae, green algae, lobose and filose testate amoebae, and, possibly, fungi, ciliates, and dinoflagellates, had appeared by Uinta Mountain Group time (Porter, 2004, and references therein). Biological and ecological complexity was also increasing: complex multicellularity, biomineralization, and sex had all been invented, and multi-tiered food webs had begun to appear (Butterfield, 2000; Porter and Knoll, 2000; Porter et al., 2003). At the end of the Neoproterozoic, this diversification culminated in a remarkable radiation of both animals and protists, commonly known as the “Cambrian explosion.” The Uinta Mountain Group records little of these events, however. Although fossils are moderately to well preserved throughout the succession, assemblages found thus far are limited in both diversity and morphological complexity. Most beds record simple, smooth-walled microfossils collectively grouped under the genus Leiosphaeridia and/or filamentous microfossils probably representing the sheaths of bacteria. Other beds preserve monospecific blooms of the bacterial aggregate, Bavlinella faveolata. More complex eukaryotic fossils, including vaseshaped microfossils (VSMs) and ornamented acritarchs, occur in the Uinta Mountain Group but are relatively rare (e.g., two out of 31 fossiliferous samples examined by Nagy and Porter [2005]

Neoproterozoic Uinta Mountain Group of northeastern Utah

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(2005). Sampling focused on the undivided strata; like the western Uinta Mountain Group, most samples yielded simple filaments, Leiosphaeridia sp., and Bavlinella faveolata. Only one specimen, from the Leidy Peak locality (Day 1, Stop 4), yielded more complex fossils, including ornamented acritarchs and possible VSMs. A single sample from the Jesse Ewing Canyon Formation yielded filaments and Leiosphaeridia sp. CORRELATION

Figure 9. Photo of the formation of Outlaw Trail in the newly subdivided Uinta Mountain Group in the eastern Uinta range. This is a 50–70-m-thick unit (thickening to >100 m to the west) and is finegrained, organic-rich, fossiliferous, and a contains a distinctive set of sedimentary structures that suggest deposition along a shoreline.

had these “complex” fossils). The limited presence of eukaryotes in the Uinta Mountain Group cannot be explained by either preservation or by age; the coeval Chuar Group, for example, records comparably preserved, but much more diverse, fossils. Instead, it is likely that the eukaryotes were excluded from the Uinta Mountain Group due to unfavorable environmental conditions. Western Uinta Mountain Group Paleontology Much more paleontological data exist for the western part of the Uinta Mountain Group. All early paleontological work on the unit (Hofmann, 1977; Nyberg, et al., 1980; Nyberg, 1982a, 1982b; Vidal and Ford, 1985) is limited to these strata, and the majority of the samples from recent work (Nagy and Porter, 2005; Dehler et al., 2006) come from here. The Red Pine Shale is the most thoroughly studied; it has yielded a variety of filaments, Bavlinella faveolata, Leiosphaeridia sp., ornamented acritarchs, vase-shaped microfossils, and the macroscopic carbonaceous compression fossil, Chuaria circularis. Though not as well studied, the Mount Watson Formation and the formation of Moosehorn Lake are also fossiliferous. The former has yielded Leiosphaeridia sp. and the ornamented acritarch Trachysphaeridium laufeldi (Vidal and Ford, 1985), and the latter has yielded filaments and Leiosphaeridia sp. (Nyberg, 1982a; Nagy and Porter, 2005). VSMs may also be present in the formation of Moosehorn Lake (Nyberg, 1982b, Dehler et al., 2006), but this has yet to be confirmed. Eastern Uinta Mountain Group Paleontology The eastern Uinta Mountain Group has been studied only recently, by Nagy and Porter (2005) and Sprinkel and Waanders

It is unclear how the western and eastern Uinta Mountain Group correlate because of the lack of detailed stratigraphic information in the central and eastern parts of the range. Research efforts on these key areas of the Uinta Mountain Group are underway in the form of mapping, measuring section, shale geochemistry, biostratigraphy, and C-isotope stratigraphy (Dehler et al., 2006), and a preliminary correlation chart is shown in Figure 3. The most complete stratigraphy is available in the eastern Uinta range where the base and the eroded top of the Uinta Mountain Group are exposed. The <770 Ma formation of Outlaw Trail, in the lower-middle part of the eastern Uinta Mountain Group, may be an eastern correlative of the shale-rich Dead Horse and Mount Watson formations, since it is one of the only significant shale intervals in the eastern Uinta Mountain Group and could correspond with the most shale-rich part of the western Uinta Mountain Group. It is hypothesized that the eastern Uinta Mountain Group section represents older and correlative strata to the majority of the western Uinta Mountain Group, except there is no lithostratigraphic equivalent of the Red Pine Shale, and there are mapping relationships that suggest that the Red Pine Shale is truncated eastward (Sprinkel, 2002). Regional correlation has long been suggested between the Uinta Mountain Group, the Big Cottonwood Formation, and the Pahrump Group and Chuar groups (e.g., Link et al., 1993). Until recently, however, there were no robust geochronological constraints on any of these units and microfossil information was limited (e.g., Crittenden and Peterman, 1975; Vidal and Ford, 1985). It is now known that the Chuar Group is 742 to ca. 770 Ma, the majority of the Uinta Mountain Group is younger than 770, and the upper Uinta Mountain Group is very similar to the 742 Ma upper part of the Chuar Group (Fig. 10) (Karlstrom et al., 2000; Williams et al., 2003; Fanning and Dehler, 2005; Dehler et al., 2006). The geochemical and petrographic similarities between the Uinta Mountain Group and the Big Cottonwood Formation to the west in the Wasatch Range strongly suggest that these units are related (Condie et al., 2001), as is the fact that the Big Cottonwood Formation is part of the same reactivated Laramide structure (the Uinta anticline or Uinta arch), suggesting that these units may have formed in the same basin along the same Neoproterozoic structure (Fig. 1). Similar fossils and Cisotope variability are observed in the Beck Springs Dolomite of the middle Pahrump Group (Prave, 1999; Corsetti and Kaufman, 2003), suggesting that it is also the same age (Fig. 10) (Dehler

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Figure 10. Regional correlation between the Uinta Mountain Group and the Pahrump Group of Death Valley, the Chuar Group of Grand Canyon, Arizona, and the Big Cottonwood Formation of the Wasatch Range, northern Utah. The recent dates from the Chuar and Uinta Mountain groups, in combination with robust correlation of C-isotope and fossil data sets or provenance data, strongly suggest that these units were deposited at about the same time and may even be part of the same regional seaway (Karlstrom et al., 2000; Dehler et al., 2001a, 2001b; Williams et al., 2003; Fanning and Dehler, 2005). Correlation with the Big Cottonwood Formation is based on sandstone compositional similarities (e.g., Condie et al., 2001). Correlation with the middle Pahrump Group is based on C-isotope and fossil data sets (Dehler et al., 2001b). Pahrump Group: BSD + “tb”—Beck Springs Dolomite and the transitional beds; K Pk—Kingston Peak Formation. Chuar Group: NF—Nankoweap Formation; Kw. Fm.—Kwagunt Formation; SF—Sixtymile Formation. Big Cottonwood Formation: LWF—Little Willow Formation; fML—formation of Moosehorn Lake. Modified from Dehler et al. (2001a).

et al., 2001a, 2001b). All of these units are, in part, marine, and all but the Beck Springs Formation record intracratonic rifting (Link et al., 1993; Timmons et al., 2001; Condie et al., 2001; Prave, 2004, personal commun.). These correlations indicate a rift-related, intracratonic seaway flooding a good part of the western Laurentia margin at ca. 750 Ma (Chuar-Uinta Mountain Pahrump groups [ChUMP] hypothesis of Dehler et al., 2001b). These correlations would make the Mineral Fork Formation, stratigraphically above the Big Cottonwood Formation, younger than ca. 740 Ma (Fig. 10). This is consistent with new U-Pb ages from zircons in the Pocatello Formation, which indicate a 709– 667 Ma age (Fanning and Link, 2004). The Pocatello Formation has long been correlated with the Mineral Fork Formation (e.g., Link et al., 1993). PALEOGEOGRAPHY The Uinta Mountain Group is most commonly interpreted as a large braided fluvial system with a trunk stream flowing westward and tributary streams flowing southward (e.g., Sanderson, 1984; Condie et al., 2001). The streams are flowing within an

east-west–trending structural trough and ultimately meeting the sea, somewhere around the Wasatch Range to the west (Fig. 11) (represented by the Big Cottonwood Formation) (Sanderson, 1984; Sanderson and Wiley, 1986; Ehlers and Chan, 1999; Condie et al., 2001). In contrast, Wallace (1972) interpreted the western Uinta Mountain Group as part of a fluvio-marine basin, with the shoreline coincident with the modern east-west divide, and the basin margin extending beyond the south flank of the modern range (Fig. 11). Results from the Red Pine Shale, and similar mudstones eastward in the Uinta Mountain Group, also suggest marine-deltaic deposition in a seaway that extended as far east as Flaming Gorge, and maybe, at times, to the Colorado border (formation of Outlaw Trail?). This seaway could have been of regional scale and extended as far southward as the Grand Canyon and Death Valley areas (“ChUMP seaway”) (Link et al., 1993; Dehler et al., 2001a, 2006). BASIN TYPE AND TECTONIC SETTING The most current views on the tectonic setting of the Uinta Mountain Group are that it was an east-west–trending intracra-

Neoproterozoic Uinta Mountain Group of northeastern Utah

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Figure 11. (A) Paleogeographic interpretation of the Uinta Mountain Group by Wallace and Crittenden (1969). Note: Interpreted strandline roughly coincides with east-west divide of Uinta Mountains. (B) Paleogeographic and tectonic interpretation by Condie et al. (2001). Work by Ball and Farmer (1998) suggests that the quartz arenites and shales found close to the axis of the range (“strandline” of Wallace and Crittenden [1969]) are derived from eastward Proterozoic cratonic sources as opposed to being reworked sands from the Archean source to the north. Arrows denote generalized paleocurrent directions.

tonic rift occupying roughly the same area as the modern Uinta Mountain range (Fig. 11) (e.g., Ball and Farmer, 1998; Condie et al., 2001). Facies, petrographic, and geochemical analyses suggest that the basin was bounded by an active fault on the northern edge, but the structural setting of the other basin margins remains unknown (Wallace, 1972; Sanderson, 1984; Condie et al., 2001). Interestingly, there are no paleocurrent data indicating northward flow, just southerly and westerly flow directions (Wallace, 1972; Sanderson, 1978; Condie et al., 2001). This suggests that the

southern basin margin was extremely low relief and/or farther away than the southern modern range boundary. Seismic profiles and mapping in the eastern Uinta Mountains indicate that the Uinta Mountain Group may thin from the north (~7 km thick) to the south (4–5 km thick) (Hansen, 1965; Stone, 1993; De Grey, 2005). The thickest part of the Uinta Mountain Group is coincident with the Uinta-Sparks fault zone on the north flank and suggests that this and/or related structures were active normal faults during Uinta Mountain Group deposition.

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Many workers have suggested that the Uinta Mountain Group was deposited in an aulacogen (e.g., Sears et al., 1982; Stone, 1993), but this is unlikely as it overlies cratonic rocks, is not associated with oceanic crust, and is ~500 km from the edge of the Proterozoic craton edge (Condie et al., 2001). Furthermore, this interpretation requires the presence of the Cordilleran (or some earlier) passive margin at Uinta Mountain Group time, which did not develop until infra-Cambrian time (Bond and Komintz, 1984; Colpron et al., 2002). Considering regional correlations with the Big Cottonwood Formation, and the Pahrump and Chuar groups (Fig. 10), there appears to have been a phase of intracratonic rifting and marine flooding at ca. 750 Ma. This is significantly earlier than the two phases of rifting suggested by Prave (1999) of 700 Ma and 600 Ma. Therefore, it is hypothesized that there were at least three phases of rifting associated with the protracted breakup of the western Laurentian margin. FIELD TRIP DAY 1—SALT LAKE CITY TO RED CANYON LODGE NEAR FLAMING GORGE Travel on Day 1 will be from Salt Lake City over the Mirror Lake Highway Scenic Byway to Red Canyon Lodge (Figs. 1 and 12). This segment of the trip offers opportunities to view and discuss the stratigraphy, sedimentology, paleontology, and geochemistry of the Uinta Mountain Group in the western and central Uinta Mountains. Road Log to Stop 1, Day 1 Cumulative Mileage 0.0 4.7 10.5 29.4 33.1 33.4 45.2 45.4

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Directions Start from the Salt Lake City convention center at 7:30 a.m., and take I-15 south. Intersection of I-80 and I-15; take I-80 east toward Cheyenne, Wyoming. Cross Wasatch fault (Bryant, 1990). Take U.S. Hwy 40 south to Heber City and Vernal. Take Exit 2 to Park City and Kamas. Turn left at the bottom of the exit and travel east on State Road (SR) 248. Intersection of SR 248 and SR 32. Turn left on SR 32. Intersection of SR 32 and the Mirror Lake Scenic Byway (SR 150). Turn right on SR 150 to Mirror Lake. The Mississippian Madison Limestone is exposed on both sides of the road in this area. This unit rests unconformably on the Uinta Mountain Group in this area. Pull over on a large dirt turnout on the right side of the road.

Stop 1—Introduction to the Western Uinta Mountain Group Stratigraphy The Neoproterozoic Red Pine Shale of the Uinta Mountain Group is exposed along the creek on the right. This is the first glimpse coming from the west of an outcrop of the Uinta Mountain Group. The Red Pine Shale is the uppermost unit in the western Uinta Mountain Group (Figs. 2 and 3). As the field trip continues over SR 150 to the east, the older strata are exposed as the road cuts toward the core of the Uinta anticline. The next lowermost unit exposed along the road is the red to pink formation of Hades Pass, and below that the white Mount Watson Formation and green to white basal strata can be seen along the road in the highest part of the range. The western Uinta Mountain Group is subdivided by sandstone composition and lithology (sandstone versus shale) (Fig. 2). There is a very distinct pattern across the range; the units are more arkosic on the north flank and more quartz rich on the south flank (Sanderson, 1978; Wallace, 1972). These compositional differences, plus the presence of shales and key sedimentary structures, show facies changes in a north-south and an east-west pattern. All of these units appear to be gradational between one another, and all facies appear to be genetically related (Fig. 3). Interpretations are variable for these western units. Sanderson (1984) interpreted the whole Uinta Mountain Group to represent a braided system, and Wallace (1972) viewed the strata to indicate a dynamic fluvio-deltaic marine system. In the next two days, many of these units will be visited to critically evaluate these paleoenvironmental interpretations. This is a very typical outcrop of the Red Pine Shale. It is predominantly shale with subordinate thin sandstone and siltstone beds. The shale is organic-rich and fossiliferous. The sandstone and siltstone beds are lenticular to tabular and exhibit a range of sedimentary structures suggesting deposition in a deltaic system. The facies exposed here represents a pro-delta environment. Road Log to Stop 2, Day 1 Cumulative Mileage 59.4 67.0

Directions The formation of Hades Pass crops out on a slope on either side of SR 150. Murdock Basin Road intersection is on the right. There is a good view of Bald Mountain (11,943 ft) to the north. The Mount Watson Formation is exposed on the left and is on the footwall block of the Hoyt Canyon normal fault (Laramide age). To the south, the formation of Hades Pass forms the cliffs, and the Hoyt Canyon fault is between here and the Hades Pass cliffs. The formation of Hades Pass (1825–3600 m thick) varies from quartz arenite

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Dehler et al. to arkosic arenite and has some significant shale intervals (tens of meters thick). This unit was interpreted by Wallace (1972) to represent a probable fluvial origin; however, considering reported bidirectional paleoflow indicators (Wallace, 1972) and the presence of a significant amount of mud during deposition, alternate depositional environments, such as a tidal area and/or delta plain would be more likely. Slate Gorge overlook. Interbedded shale and tabular siltstone and sandstone beds in the formation of Moosehorn Lake can be observed here. Bryant (1992) mapped the Hoyt Canyon fault south of here (near Stop 1), where the formation of Hades Pass to the south is juxtaposed against the formation of Moosehorn Lake and Mount Watson Formation to the north. Turn left (N) into the Bald Mountain photo viewpoint turnout and parking lot. It is recommended to view the Mount Watson Formation here and then carefully walk along the left side of the road 0.2 mi to the Bald Mountain picnic area, where the formation of Moosehorn Lake can be visited. Take the trail toward Bald Mountain to view the outcrop character of the Mount Watson Formation.

Stop 2—Basal and Middle Western Uinta Mountain Group The Mount Watson Formation (≤1000 m thick) comprises light-colored orthoquartzite, quartz arenite, and subarkosic arenite (Figs. 4 and 5). Sedimentary structures exposed at this locality are thin- to medium-bedded trough crossbeds, soft-sediment deformation, and symmetric and interference ripplemarks. Paleoflow direction at this stop is dominantly to the northwest and symmetric ripplecrest orientations indicate a locally northsouth–trending shoreline. Other sedimentary features observed by Wallace (1972) are distinctive 15–20-m-thick sets of thickly bedded planar tabular foresets, trough crossbeds tens of meters wide and long in plan view, and thin shale beds. There are subordinate intervals of green and red shale in this unit, which thicken to the east, where the unit is called the formation of Dead Horse Pass. This shaley correlative may correlate with a green shale unit, the formation of Outlaw Trail, in the eastern part of the range (Figs. 3 and 9) (Dehler and Sprinkel, 2005). Northward and northwestward, the Mount Watson Formation grades into the formation of Red Castle, which is dominantly arkosic arenite with fewer interbedded shale intervals (Figs. 2 and 3). There are two different environmental interpretations for the Mount Watson Formation and equivalent units. Sanderson (1978, 1984) interpreted these strata to represent a braided sheetwash plain and extended this interpretation to the entire Uinta Mountain Group. Wallace (1972) interpreted this formation to be a dynamic system with a fluvial component, as well as facies representing

shoreline, deltaic, and offshore deposition. Wallace’s interpretation describes tributary streams carrying arkosic sediment from the north, with reworking of local stream deposits into quartz arenite and subarkosic sediment along an east-west–trending shoreline coincident with the modern Uinta Mountain divide. He specifically recognized deltas, tidal flats, mudflats, and offshore bars. Sanderson’s interpretation describes tributary streams carrying arkosic sediment from the north, feeding a larger braided system from the east carrying quartzose sediment. Provenance studies by Ball and Farmer (1998) indicate that the quartz arenite source was derived from a chemically different source than the arkosic source, indicating that the quartz arenite was not derived through mechanical reduction of arkosic sediment. The formation of Moosehorn Lake (<360 m) comprises green to gray to yellow shale, with subordinate thin to medium beds of arkosic siltstone and sandstone (Fig. 4). Sedimentary structures that can be viewed here are thin to medium bedding and soft sediment deformation. Wallace (1972) interpreted this unit to represent tidal flats, submerged to emergent delta plains, and/or lagoonal environments, with the interbedded sandstone beds indicating delta front or offshore bar deposition. Only simple fossils have been found in this formation: filaments and smooth-walled microfossils of the genus Leiosphaeridia (Fig. 13A–13C and 13F; Nyberg, 1982a; Nagy and Porter, 2005). The former likely represent the empty sheaths of filamentous bacteria; the affinities of the latter, often referred to as “leiosphaerids,” are more problematic because the genus likely contains a range of unrelated taxa. Some specimens may represent prokaryotes; others have been interpreted as the resting stages of simple eukaryotic algae. Nyberg (1982a) also reported VSMs from this unit, but the two specimens he illustrated have irregular morphologies more reminiscent of mineral grains than of the rounded tests of VSMs. Specimens representing “possible VSMs” are reported by Dehler et al. (2006); their preservation is not good enough to allow confident attribution. The axis of the Uinta anticline was mapped by Bryant (1992) near here. Bryant (1992) also mapped faults on Murdock Mountain to the southeast that place Mount Watson Formation next to the formation of Moosehorn Lake.

Figure 13. Examples of fossil species from the Uinta Mountain Group. (A–C) Leiosphaeridia sp.: (A) a solitary specimen of Leiosphaeridia; (B) a Leiosphaeridia specimen exhibiting a medial split, a likely excystment structure; (C) aggregates of Leiosphaeridia cells. (D–E) Bavlinella faveolata. (F–H) Filaments, both unbranched (F) and branched (G–H). (I)—Satka colonialica. (J–M) Examples of “complex” acritarchs found in the Uinta Mountain Group: (J) Trachysphaeridium laminaritum; (K) a species with an outer envelope around an inner spheroid; (L) a species with ornamented walls; (M) a species covered in short spines. (N) Two vase-shaped microfossils, attached at their apertures. (O) The modern testate amoeba, Difflugia lucida, undergoing asexual reproduction. (I) and (J) are from the Chuar Group; they are representative of the same species found in the Uinta Mountain Group. (O) is courtesy of Ralf Meisterfeld.

Neoproterozoic Uinta Mountain Group of northeastern Utah

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Dehler et al.

16 Road Log to Stop 3, Day 1

199.6

Cumulative Mileage

201.3

74.8

74.9

75.6

77.5 79.2 82.8 96.9 121.1 122.9

152.2 157.4 160.4

182.1 190.5 192.9

Directions Bald Mountain Picnic area is on the left. From here, a short hike accesses the basal formation of Moosehorn Lake and the overlying Mount Watson Formation. After Stop 2, leave Bald Mountain parking, turn left on SR 150, and continue north to Mirror Lake and Evanston, Wyoming. Duchesne County line. Outcrop of quartz arenite of the Mount Watson Formation is on the right and faulted formation of Moosehorn Lake on Murdock Mountain is ahead and to the right. Just before turnoff to Hayden Peak overlook, the contact between the formation of Moosehorn Lake and the overlying Mount Watson Formation can be viewed on the left. Hayden Peak and Moosehorn Lake overlook (alternative Stop 2). Hayden Peak (12,479 ft) is in full view to the northeast. Hayden Peak is capped by the formation of Hades Pass (reddish rocks) and is underlain by the light-colored Mount Watson Formation. Faults cut the rocks north and south of Hayden Peak. To the southeast, the green shale and sandstone beds are the formation of Moosehorn Lake. Pass Lake. View of formation of Hades Pass in cliffs the on peaks to the north. Fine-grained interbeds of Mount Watson Formation. Formation of Hades Pass is exposed in road cut on the left. The road crosses the approximate position of the North Flank fault. Evanston city limits. Continue through outer Evanston to I-80 eastbound onramp. Turn right onto I-80 east toward Rocks Springs and Cheyenne. Travel east to the Fort Bridger exit. Take the Fort Bridger exit and travel east in the I-80 business loop through Fort Bridger. Urie, Wyoming. Turn right at intersection and travel south to Mountain View, Wyoming. Mountain View, Wyoming. Stay on Wyoming SR 414. The road near the south end of Mountain View will split; keep left to stay on SR 414 to Lone Tree, Wyoming. Lonetree, Wyoming. Burnt Fork, Wyoming Spirit Lake Road is on the right; turn right on Spirit Lake Road, and travel south. To the east (left) is Phil Pico Mountain.

201.8

206.1 207.2

Pennsylvanian-Permian Weber Sandstone is exposed in road cuts. We continue to go down section as we travel up the road in Birch Canyon. Madison Limestone forms the cliff to the east and west of the road. Just south of here is the angular unconformity between the underlying Red Pine Shale and the overlying Madison Limestone. There are no Cambrian strata exposed in this area. Red Pine Shale exposed in road cuts. Samples collected here have yielded Leiosphaeridia spp., filaments, Satka colonialica, and Trachysphaeridium laminaritum (Sprinkel and Waanders, 2005; Nagy and Porter, 2005). Satka colonialica is characterized by spherical to elongate envelopes with impressions of cells that were once inside (Fig. 13I). Trachysphaeridium laminaritum is an organic-walled microfossil with small, tightly packed, round depressions covering its surface (Fig. 13J). Many specimens have a regular circular opening, likely representing an excystment structure; i.e., an “escape hatch” through which a cell can escape the resistant outer wall (the “cyst”) that has protected it during its resting period. Both species are known from early to mid-Neoproterozoic successions, including the Chuar Group, Arizona, and the Visingsö Beds, Sweden (Vidal and Ford, 1985). Spirit Lake Lodge Road on left; stay on the U.S. Forest Service road (USFS 221). Turn left at the intersection of USFS 221 and the access road to the borrow pit in the Red Pine Shale. After Stop 3, continue east on the USFS 221.

Stop 3—Last Gasp of the Red Pine Shale? Just west of here by ~10 km, in the Hoop Lake area, there are thick exposures (>500 m) of Red Pine Shale comprising interbedded arkosic sandstone, siltstone, and organic-rich shale. No Red Pine Shale has been formally mapped in this area, but it is clearly still present (Fig. 1). We are not far from the last exposure of the Red Pine Shale on the north flank. It is unclear to us at this time if the Red Pine Shale undergoes a facies change from predominantly shale to sandstone with shale interbeds, or if the Red Pine Shale has been removed by erosion or faulted out and the sandstone and shale interbeds are the formation of Hades Pass stratigraphically downsection. On the south flank of the Uinta Mountains, the Red Pine Shale is exposed from the western end of the range to the east side of Ashley Creek (Fig. 1). There, the Red Pine Shale is truncated and in angular discordance with the Mississippian Madison Limestone. Preliminary δ13Corg data from the Henry’s Fork section, ~25 km west of here, shows values from −19.6 to −26.5‰, and

Neoproterozoic Uinta Mountain Group of northeastern Utah the Ashley Creek section, ~40 km south of here, values range from −17.1 to −27.7‰ (Fig. 2). Values from these localities are similar to the values in the composite δ13Corg curve generated from western shale localities (to be discussed at Stop 6, Day 2) and are encouraging for potential use in correlation of the Red Pine sections across the range (Fig. 2). Hoop Lake paleontology samples have yielded the same simple fossil assemblages seen elsewhere: filaments, including some branching specimens (Fig. 13G and 13H), Leiosphaeridia sp., and Bavlinella faveolata. This last fossil is usually found by itself, and is thought to have been an opportunistic taxon, creating blooms in environments adverse to other taxa (Knoll et al., 1981). Here B. faveolata is found in association with filaments and Leiosphaeridia sp. Road Log to Stop 4, Day 1 Cumulative Mileage 208.8

216.7

217.5 217.8

220.2

17

turn around and retrace route (USFS 218) to get to SR 44. Stop 4—The North Flank Fault and the Uinta Mountain Group The view is to the west and north: the Uinta Mountain Group consists of red sandstone, siltstone, and mudstone that form the bowl of an amphitheater. The rim of the amphitheater is formed by the gray cliff of the Mississippian Madison Limestone, which is locally referred to as “The Palisades.” The southwestern branch of the Uinta fault zone placed the Uinta Mountain Group (on the south) next to the Madison Limestone (Fig. 14). West of the Sheep Creek area, the fault likely continues along the base of the Madison cliff to west of Long Park Reservoir, where it may cut down into the Uinta Mountain Group and place the prob-

Directions Intersection of USFS 221 and access road USFS 014 to Long Park Reservoir. The Uinta Mountain Group along the Long Park Reservoir access road comprise mostly medium to thick beds of red quartz sandstone interbedded with greengray shale beds. The general dip of the Uinta Mountain Group is ~12°N; however, near here the Uinta Mountain Group is folded and steeply tilted. These sandstone and interbedded shale beds are similar to the succession of rocks seen at Sheep Creek Canyon (our next stop). Continue on USFS 221. Intersection of USFS 221 and the Sheep Creek Canyon Geological Area loop road. Turn left onto the loop road (USFS 218) to the Sheep Creek Canyon Geological Area. Intersection with USFS 93 (Death Valley Road). Continue west on USFS 218. Beginning of switchback. View to the westnorthwest of Sheep Creek and Mahogany Draw in the Sheep Creek Canyon Geological Area (Schell, 1969; Sprinkel et al., 2003). The deep red strata are the sandstone beds of the Uinta Mountain Group. Interbedded with the sandstone beds are green-gray shale beds that contain microfossils. The steep hill north of the road is Windy Ridge. Windy Ridge consists of steeply dipping and overturned beds of the Mississippian Madison Limestone. The road is built on the Uinta Mountain Group. The Uinta fault zone is near the base of the Madison Limestone. Pull over in a large turnout at the boundary of Sheep Creek Canyon Geological Area and Palisades Memorial Park overlook. After Stop 4,

Figure 14. Photo looking north at the Uinta Fault zone. The Uinta Mountain Group is on the hanging wall, and the Madison Limestone (cliff at top) is on the footwall of a Laramide-age reverse fault dipping to the south. This fault is one of several that bound the northern Uinta range. The Uinta Mountain Group stratigraphy undergoes character changes in this part of the range and this is therefore a key area for understanding how the eastern and western Uinta Mountain Group stratigraphy correlate across the range. It is still unclear if and how structure played a role in the stratigraphic changes. Photo by Andy Brehm.

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able upper-middle part of the Uinta Mountain Group over the Red Pine Shale. An alternative interpretation is that interbedded sandstone and shale exposed here is equivalent to the Red Pine Shale to the west, representing a facies change, which would not require the thrust fault to continue to the west. Green-gray shale beds in this area mostly contain the simple fossils seen elsewhere: Leiosphaeridia sp. and filaments, some of them branching (Fig. 13G and 13H). A single sample from this locality, however, preserves a more diverse assemblage of relatively complex acritarchs, typical of Neoproterozoic shallow water environments (Butterfield and Chandler, 1992). These include species that possess an outer envelope around an inner spheroid (Fig. 13), species with ornamented walls (Fig. 13J and 13L), and species with a diversity of spines or “processes” (Fig. 13M). “Probable VSMs” have also been reported from these samples, but their poor preservation prevents confident assignment (Dehler et al., 2006). Bedding in the Sheep Creek Canyon Geological area, and beyond, forms the north flank of the Uinta arch, a broad asymmetrical anticline ~30 mi wide and 150 mi long (Fig. 1) (Hansen, 1965; Sprinkel, 2003). The Uinta arch consists of two large domes that are aligned east-west and are separated by a shallow structural saddle, which is crossed by U.S. Hwy 191–Utah Hwy 44 from Vernal to Manila (Hansen, 1965; Sprinkel, 2003). Thus, the Sheep Creek Canyon Geological Area is on the eastern part of the western dome. Road Log to Stop 5, Day 1 Cumulative Mileage 223.7

226.7

227.5 237.9

240.4

Directions Intersection of Spirit Lake Road (USFS 221) and Deep Creek Road (USFS 539). Continue east on USFS 218, which is on the Uinta Mountain Group. Intersection of Sheep Creek Canyon Geological Area loop road (USFS 218) and SR 44. Turn right on SR 44 and travel east to Greens Lake Recreation area and the Red Canyon overlook. Dowd Mountain turn off; stay on SR 44. Intersection of Green Lakes Recreation area and Red Canyon overlook: turn left on USFS 095 and travel north. Either drive to the Green Lakes Recreation area, which includes the Red Canyon overlook, or park at the Red Canyon Lodge and walk the short rim trail to the overlook. After this stop, stay overnight at Red Canyon Lodge.

where Hansen (1965) puts a structural trough between two domes that make up the Uinta anticline. The strata take on a different look in this area, and it could be that the domes were once subbasins that, in part, controlled deposition of the Uinta Mountain Group. As will be shown on Day 2, the eastern Uinta Mountain Group appears to represent a fluvial-dominated system, whereas in the western Uintas, the strata reflect marinedeltaic, as well as fluvial, deposition. The Uinta Mountain Group seen from the overlook consists of interbedded sandstone and red shale beds. Sedimentary structures in the local vicinity are trough crossbeds and low-angle crossbedding, similar to much of what we will see on Day 2. Sedimentological and stratigraphic analyses have yet to be conducted in this area to determine the depositional environment and local correlation. The strata in this area are particularly indistinct and will be a challenge to correlate even locally. These strata are likely equivalent to the middle-upper part of the western Uinta Mountain Group (formation of Hades Pass or Red Pine Shale) (Fig. 3). FIELD TRIP DAY 2—RED CANYON LODGE, FLAMING GORGE AREA, OVER DIAMOND PLATEAU, AND RETURN TO SALT LAKE CITY Travels on Day 2 are from Red Canyon Lodge over Flaming Gorge, down Jesse Ewing Canyon, across Browns Park, over the Diamond Mountain Plateau, through Vernal, up the north Fork of the Duchesne River, and back to Salt Lake City (Fig. 12). Stops will exemplify the different formations and proposed subdivisions of the eastern Uinta Mountain Group, and we will discuss correlation between the Uinta Mountain Group and other strata regionally (Figs. 3 and 10). We will also examine different facies and facies associations and discuss interpretation of paleogeography and paleoclimate based on sedimentology, paleontology, stratigraphy, geochemistry, and geochronology. Road Log to Stop 1, Day 2 Cumulative Mileage 0.0

3.7

12.1 Stop 5—East Meets West: Character Changes in the Uinta Mountain Group 14.3 This is a profound area for trying to understand the Uinta Mountain Group stratigraphy (Fig. 3). This overlook sits about

Directions Leave Red Canyon Lodge and travel south to the intersection with SR 44. Turn left on SR 44 to Flaming Gorge Dam and Vernal. Intersection of SR 44 and U.S. Hwy 191. Turn left on Hwy 191, and travel north to Flaming Gorge Dam. Dutch John, Utah. This town was built by the Bureau of Reclamation to house workers constructing Flaming Gorge Dam. Pull over along the side of the road on the right, adjacent to the outcrop of sandstone and green shale on the right (east) side of the road.

Neoproterozoic Uinta Mountain Group of northeastern Utah Stop 1—Organic-Rich Shale Beds in the Uinta Mountain Group This outcrop exhibits a significant amount of variability in grain size and color. At the southern end of the outcrop, there are many shale intervals (meters thick) that are green to gray, indicating the presence of organic matter. Interbedded with these shale beds are thin to medium beds of green to red sandstone, pebbly sandstone, and local conglomerate. The pebble-sized clasts are angular to rounded and are composed of quartzite derived from the underlying Red Creek Quartzite. These shale beds contain the same simple assemblages seen elsewhere: filaments and Leiosphaeridia sp., some of the latter as solitary cells (Fig. 13A and 13B), some in aggregates (Fig. 13C). Also present is the ornamented acritarch Trachysphaeridium laminaritum (Fig. 13J; Sprinkel and Waanders, 2005; Nagy and Porter, 2005). Road Log to Stop 2, Day 2 Cumulative Mileage 15.3

21.0

31.0

33.6 39.0

Directions The Uinta-Sparks fault zone, dipping to the south, is a reverse fault that places the Neoproterozoic Uinta Mountain Group over the Triassic Chinle Formation and Jurassic Nugget Sandstone (Sprinkel, 2002). Intersection of U.S. Hwy 191 and the Clay Basin road. Turn right (east) on the Clay Basin road. The road parallels the Cretaceous Mesaverde Group on right and the Cretaceous Fort Union Formation on left. The road travels mostly on the lower part of the Fort Union Formation but drops onto the Mesaverde Group just west of Clay Basin. There, the UintaSparks fault zone is exposed and places the Archean-Paleoproterozoic Red Creek Quartzite over the Mesaverde Group. Clay Basin and the Clay Basin gas field. Clay Basin is floored by the Baxter Shale and surrounded on the north by cliffs of the Cretaceous Mesaverde Group and Tertiary Wasatch Formation. The hills to the south are composed of Red Creek Quartzite and the Uinta Mountain Group. The Uinta-Sparks fault zone is near the base of the hills to the south. Intersection; turn right and continue through Clay Basin to Browns Park. The Uinta-Sparks fault zone here places the Uinta Mountain Group over the Cretaceous Baxter Shale; however, the zone is several hundred feet wide here. Continue on Browns Park Scenic Byway road and descend down

40.7

19

Jesse Ewing Canyon to Browns Park. The Jesse Ewing Canyon Formation and undifferentiated Uinta Mountain Group are exposed in the hills on both sides of the road. Pull over along the right side of the road where an unnamed fault places the Neoproterozoic Uinta Mountain Group down to the south next to the Archean-Paleoproterozoic Red Creek Quartzite. It is best to park up the road from this fault just a few hundred meters and walk down along the road for the next ~0.5 mi to look at the spectacular exposures of this basal unit. Watch for vehicles!

Stop 2—The Base of the Uinta Mountain Group On the footwall of this unnamed normal fault is the Red Creek Quartzite, a likely correlative with the Farmington Canyon Complex in the Wasatch Range. This Paleoproterozoic-Archean (?) unit has several metamorphic facies that were mapped by Hansen (1965). On the hanging wall of this fault is some of the best exposure of the Jesse Ewing Canyon Formation, the basal unit of the eastern Uinta Mountain Group (Fig. 2) (Sanderson and Wiley; 1986). Although it is a clast-supported cobble breccia at this spot, laterally and vertically, this unit varies radically and includes beds of cobble conglomerate, pebble conglomerate and breccia, sandstone, siltstone, and significant red to green to gray shale intervals (Fig. 7) (Sanderson and Wiley, 1986). These relationships are nicely exposed down and along the road for the next 0.5 mi. The coarser-grained facies become subordinate within this 0.5 mi distance, and shale becomes dominant. Although red in appearance, much of this shale is gray to black on the inside and contains microfossils including leosphaerids and carbonaceous filaments. Sedimentary structures in the Jesse Ewing Canyon Formation are abundant. In the shale intervals, parallel laminations, ripplemarks (symmetric and asymmetric, ladder), soft-sediment deformation, mudcracks, evaporite pseudomorphs, potential bipolar crossbeds, tool marks, mud chips, load casts, wavy bedding, and hummocky-cross stratification have been observed. In the coarse-grained intervals, trough and tabular crossbedding, fining- and coarsening-upward sequences, and weak imbrication are common. Sanderson and Wiley interpreted this unit to represent an alluvial fan complex. We suggest, based upon the lateral facies changes, the high organic content, and the presence of microfossils, that the depositional environment was, in part, subaqueous. It could be that many of the breccia and conglomerate beds also represent subaqueous deposition in the form of coarse-grained turbidites and mass-flow deposits. Recent mapping and stratigraphic studies imply that the Jesse Ewing Canyon Formation exhibits differential subsidence along an active (normal?) fault. Measured sections on the north side of this fault are ~200 m thick (Sanderson and Wiley, 1986),

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whereas sections on the south side of this fault indicate a minimum thickness of 410 m. Road Log to Stop 3, Day 2 Cumulative Mileage 41.0

41.2

42.7

44.1 44.3

Directions Sign 831; filaments and leiosphaerids were recovered from these shales (Sprinkel and Waanders, 2005; Nagy and Porter, 2005). Mountain Home fault. The Brown Park Formation is faulted down next to red shale beds of the Jesse Ewing Canyon Formation. This represents the valley bounding fault that forms the north side of the Browns Park graben. There is no Jesse Ewing Canyon Formation or crystalline basement exposed south of this structure. Correlation of the northern Uinta Mountain Group strata with the significant thickness of Uinta Mountain Group strata on the south side of the graben has not yet been successful. Intersection of Browns Park Scenic Byway and the road to the John Jarvie House. Turn right to the John Jarvie House. At the bridge, turn left on the Taylor Flat Road and cross the Green River. At the intersection, turn west (right) off Taylor Flat Road to the trailhead of Outlaw Trail.

45.8

Park the vehicles at the trailhead and hike ~0.5 mi upstream along the south side of Green River to view the formation of Outlaw Trail. At end of Stop 3, hike back to vehicles and retrace route back to the Taylor Flat Road.

Stop 3—Formation of Outlaw Trail and Subdivision of the Eastern Uinta Mountain Group along the Diamond Breaks The bedrock unit the trail follows to get to the formation of Outlaw Trail is the underlying formation of Diamond Breaks. The formation of Crouse Canyon, which overlies the formation of Outlaw Trail, is very similar in facies characteristics, composition, paleoflow direction, and depositional environment to the underlying formation of Diamond Breaks (Figs. 2, 3, 15, and 16; De Grey and Dehler, 2005). Dominant sedimentary structures in the underlying and overlying formations include trough crossbeds, low-angle crossbeds, and planar-tabular crossbeds (Fig. 16). Soft-sediment deformation, mudcracks, mudchips, and asymmetric ripplemarks are also present. Paleocurrent data from these units indicate flow to the south-southwest (Fig. 15) (De Grey, 2005). The facies associations in these sandstone formations indicate a low- to moderate-energy braid plain with higher energy main channels of a braided river (De Grey, 2005). The formation of Outlaw Trail is a mappable green to red shale unit with subordinate sandstone beds that is ~50–70+ m thick and allows the Uinta Mountain Group to be subdivided into three units in the eastern Uinta Mountains (Figs. 8 and 9) (Connor et al., 1988; De Grey, 2005; Dehler et al., 2006).

Figure 15. Rose diagrams from crossbedding in the formation of Diamond Breaks (A) and the formation of Crouse Canyon (B). The paleoflow direction is similar for both of these units and indicates predominantly southwesterly flow.

Neoproterozoic Uinta Mountain Group of northeastern Utah It is an oxidized to organic-rich shale unit with interbeds of subarkosic to arkosic siltstone and arenite (Fig. 9). Sandstone and siltstone beds are typically thin to medium bedded, yet the sandstone beds become thicker to the west, as does the whole unit. Sedimentary features include planar-tabular crossbeds, planar-horizontal laminae, symmetric, asymmetric, and interference ripplemarks, mudcracks, gypsum pseudomorphs, and soft-sediment deformation. Sandstone beds thicken westward, along with the westward thickening of the whole formation. The shale becomes more variegated and locally organic westward (Dehler et al., 2006). Filaments and leiosphaerids were recovered from these shales (Nagy and Porter, 2005). This unit

21

represents a delta plain and is likely a marine flooding event from the west. A possible local correlation is between the formation of Outlaw Trail and the Mount Watson Formation and equivalents (Fig. 3). These western units (especially the formation of Deadhorse Pass) are the thickest and most abundant shale intervals besides the uppermost Red Pine Shale. The stratigraphic position in about the middle of the succession supports this correlation. A 770 Ma detrital zircon population from sandstone in the formation of Outlaw Trail indicates that this unit can be no older than 770 Ma (Fanning and Dehler, 2005). If the local correlation with the middle western Uinta Mountain Group is

A

B

Figure 16. Two of the most common facies within the sandstone formations of the eastern Uinta Mountain Group. (A) Thick, massive tabular to lenticular sandstone beds. (B) Medium-bedded trough crossbeds. These sedimentary structures indicate deposition in a braided-river system.

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correct, then this age association can be used in the western part of the range as well. Road Log to Stop 4, Day 2 Cumulative Mileage 47.3 50.9

56.1

57.7

58.7

68.9

70.8 71.2

Directions Taylor Flat Road. Turn right to continue to Crouse Canyon Road. At the intersection with the road to Sears Canyon, stay on Taylor Flat Road to Browns Park Road. The strata exposed along the road between here and Crouse Canyon is the lowermost of the subdivided units along the Diamond Breaks, the formation of Diamond Breaks. Watch for trough crossbedding, lowangle crossbedding, and large channelforms representing a braided stream system. Swallow Canyon. The Green River cuts through a spur of Uinta Mountain Group. Swallow Canyon is not far south of the approximate projection of the eastern part of the Uinta axis (Hansen, 1965; Sprinkel, 2002). Junction of Taylor Flat and Browns Park roads. Turn right and travel south and up Crouse Canyon. Mouth of Crouse Canyon. Near the mouth of Crouse Canyon is the poorly exposed formation of Outlaw Trail. The road up Crouse Canyon travels stratigraphically upsection through the Uinta Mountain Group. The average dip of the Uinta Mountain Group from here southward is ~12°SW. There are great views of the lithofacies and sedimentary structures in the Uinta Mountain Group in the canyon. Intersection of Browns Park Road and the road to Crouse Reservoir. Turn right onto the connecting road to Crouse Reservoir. The road travels west through Uinta Mountain Group and surficial deposits. At the intersection, turn right to Crouse Reservoir. Pull over along the road. After Stop 4, turn around and travel due south to Vernal.

upper part of the eastern Uinta Mountain Group can be subdivided further. Otherwise, these sandstone strata exhibit the same facies associations as the lower sandstone formations. This east-west–trending valley is following a normal fault that was active in the Quaternary and likely the Tertiary (Hansen and Rowley, 1991). Seismic log interpretation shows a structure possibly offsetting the Uinta Mountain Group by ~<100 m (Stone, 1993). This would affect calculations of total thickness of the Uinta Mountain Group. Other structures in the area show remarkably small amounts of offset (≤10 m). These structures are oriented between ~50° and near vertical, typically trend eastwest, dip to the north or south, and are associated with multi-orientation joint sets (De Grey, 2005). Road Log to Stop 5, Day 2 Cumulative Mileage 71.6 73.3

74.4 76.6

79.9

92.4

101.9 102.4

160.9 166.8

Stop 4—Upper Eastern Uinta Mountain Group This stop is nearing the top of the upper eastern Uinta Mountain Group. Currently, we are extending the formation of Crouse Canyon all the way to the top of the group (Figs. 2 and 3). Two poorly exposed green shale beds (tens of meters thick) are mapped in this area and are similar to the formation of Outlaw Trail (Hansen and Rowley, 1991). It is not known whether these shale units are laterally continuous and therefore whether this

186.9 191.3 197.0

200.4 200.6

Directions Continue heading south. The Madison Limestone hogback with Lodore Sandstone exposed below. The Uinta Mountain Group is exposed below the Lodore Sandstone. At this intersection, veer right and continue south on Browns Park Road. At the intersection of the Browns Park Road and Jones Hole Road, turn right and travel across the Diamond Mountain Plateau. At this intersection, take the left fork to Vernal. Follow the road off the Diamond Mountain Plateau to Brush Creek. Cross Brush Creek and stay on the road to Vernal. This road will skirt around the east side of the Buckskin Hills and past the Uintah County landfill. Cross Ashley Creek and continue west to 500 East. Turn left on 500 East and travel south to U.S. Hwy 40. Turn right on U.S. Hwy 40 and travel west through Vernal. Stay on Hwy 40 through Roosevelt and to Duchesne. Duchesne. Turn right on SR 87 and travel north to the intersection with SR 35. At the intersection, turn left onto SR 35 to Tabiona, Hanna, and beyond. The road follows the Duchesne River. Tabiona. Hanna. Intersection of SR 35 and USFS 144 to upper Duchesne River drainage. Turn right on USFS 144 and travel north into the upper Duchesne River drainage toward the Hades Canyon turn off. Tintic Quartzite in road cut on right. Red Pine Shale is exposed.

Neoproterozoic Uinta Mountain Group of northeastern Utah 201.2

Pull over along the side of the road. After Stop 5, continue straight ahead.

Stop 5—Paleotopography Developed on the Red Pine Shale The uppermost unit of the western Uinta Mountain Group, the Red Pine Shale, is apparently onlapped by the Cambrian Tintic Quartzite. On the east-facing side of the canyon, the cliff of Tintic Quartzite dramatically thins to the north and pinches out over this paleohigh in the Red Pine Shale. Up the canyon, the Mississippian Madison Limestone sits directly on the Red Pine Shale. Paleotopographic highs in the Uinta Mountain Group are also exposed east of here along the Green River, just downstream from the confluence with the Yampa River. The combination of these highlands and pervasive angular unconformities in the Uinta Mountains suggest significant regional uplift after ca. 750 Ma and prior to the Middle Cambrian, likely associated with the rifting of the Laurentian margin. The Red Pine Shale in this canyon is ~1200 m thick and consists of a shale facies, an interbedded shale and siltstone and sandstone facies, and a sandstone facies (Figs. 2 and 6). The shale and siltstone and some of the sandstone are organic rich. Note the Red Pine sandstone and siltstone beds forming the low cliff along west side of Duchesne River up-canyon. Beyond that cliff, on the west side, you may be able to see the thick succession (hundreds of meters) of sandstone making up the middle and upper slope in the upper Red Pine Shale. Road Log to Stop 6, Day 2 Cumulative Mileage 203.5

207.6

Directions Road junction of Hades Canyon to Grand View trailhead to the Granddaddy Lakes area. Turn right and travel up Hades Canyon to Castle Rocks overlook near Splash dam. Pull over at the Castle Rocks overview pullout. When Stop 6 is over, turn around and retrace route down the canyon to SR 35.

Stop 6—Paleoenvironments and Paleontology of the Red Pine Shale The Castle Rocks viewpoint sits very close to an eastwest–trending normal fault, with the formation of Hades Pass on the footwall (north side) and the Red Pine Shale on the hanging wall (south side). The formation of Hades Pass will only be visible from a distance, yet is the cliffy white to reddish sandstone to quartzite with significant red shale interbeds. Where exposed, the contact between the formation of Hades Pass and the Red Pine Shale in gradational. The Red Pine Shale that is exposed to the south from the Castle Rocks overlook, on the south side of Hades Creek, is the

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site of a large ~600 m thick measured section, and two other partial sections to the west on the same ridge (Fig. 6) (Dehler et al., 2006). Sections were also measured across the Duschesne River along the lower cliff to the south and across and northward on the steep sandstone cliffs to the west (both of these sections are on the east facing slope of the main canyon). From mapping and correlating these sections, C. Dehler has calculated a thickness of 1200 m. These sections were sampled for shale geochemistry, detrital zircon analysis, paleontology, petrography, and C-isotope stratigraphy. Common sedimentary structures include parallel- to ripplelaminations, asymmetric and symmetric ripplemarks, mudcracks, normal and reverse grading, slump folds, load structures, hummocky-cross stratification, silica concretions, mud chips, rare quartzite pebbles, cut-and-fill structures, planar-tabular crossbeds, and rare associated topsets. Sandstone beds are thin to thickly bedded, and shale and siltstone intervals are typically tens to hundreds of meters thick (Fig. 6). This suite of sedimentary structures indicates prodelta, delta front, and delta plain deposition in a marine basin (Dehler et al., 2006). δ13Corg analysis of organic-rich shales reveals variability of 13.3‰ (Peedee belemnite [PDB]), with values ranging from −16.9 to −30.2 ‰ PDB (Fig. 2; Dehler et al., 2006). Preliminary total organic carbon values range from 0.07 to 5.91%. A composite δ13Corg curve in Figure 2 includes data from the two bestexposed and more completely sampled sections—the type section on the northwest flank and the Hades Creek section (Fig. 1). The variability in δ13Corg values is similar to other δ13Corg curves from middle to late Neoproterozoic marine successions (e.g., Coats Lake Group, Kaufman et al., 1997; Chuar Group, Dehler et al., 2005) suggesting that the Red Pine Shale is within this age range and is likely marine in origin. Detrital zircon data indicate that at least two sources of sand were deposited into the Red Pine basin: an arkosic Archean source from the Wyoming craton to the north and a Paleoproterozoic source from the Colorado province to the east. These data are similar to geochemical data from the Big Cottonwood Formation, suggesting another linkage between the Uinta Mountain Group and the Wasatch strata to the west (Condie et al., 2001). Like the rest of the Uinta Mountain Group, most of the samples collected from the Red Pine Shale yield simple fossils. Assemblages consisting only of filaments and leiosphaerids are common. Also common are monospecific assemblages of the bacterial aggregate Bavlinella faveolata (Fig. 13D and 13E). These are preserved in black, sulfur-rich shales, consistent with the hypothesis that B. faveolata was an anoxygenic photosynthetic bacterium that lived in euxinic environments (this would also explain the absence of eukaryotic fossils from B. faveolata assemblages; Vidal and Nystuen, 1990). The ~mm-scale carbonaceous compression fossil Chuaria circularis has also been reported from the Red Pine Shale (Hofmann, 1977; Nagy and Porter, 2005). In addition to these simple fossils, a single, silicified mudstone from the middle-upper part of the type section has yielded abundant vase-shaped microfossils (VSMs) preserved as siliceous internal molds. VSMs are tear-drop–shaped or hemispherically

Dehler et al.

24

shaped tests with a circular or polygonal aperture at one end. They have been shown to be the remains of filose and lobose testate amoebae, two groups that are common today in freshwater and terrestrial environments (Porter and Knoll, 2000). VSMs are locally abundant and widespread in rocks that just pre-date the Sturtian glaciations, but, with one possible exception, do not occur in post-Sturtian rocks. They thus provide a useful biostratigraphic marker for immediately pre-Sturtian times. Of particular interest in the Red Pine Shale population are two VSM specimens that are attached at their apertures (Fig. 13N). The same pose is found in modern testate amoebae undergoing cell division (Fig. 13O), suggesting these specimens were undergoing asexual reproduction when they died. Road Log to Salt Lake City Cumulative Mileage 218.2

252.2 264.0 268.0

Directions Intersection with SR 35; turn right on SR 35 to Kamas. The road travels up and over Wolf Creek Pass, down to Woodland, and then to Kamas. Kamas. Turn left on SR 248 to U.S. Hwy 40. SR 248 and Hwy 40. Take Hwy 40 north to I-80. Take I-80 west to Salt Lake City.

End of Day 2 and end of field trip. ACKNOWLEDGMENTS Sprinkel’s ongoing work is supported by National Cooperative Geologic Mapping Program state geologic survey mapping component (STATEMAP) funding. Dehler’s work and that of her students is supported by the National Cooperative Geologic Mapping Program university geologic mapping component (EDMAP), the Utah Geological Survey, and a New Faculty Grant from Utah State University. Porter’s work was supported by the National Science Foundation and an Academic Senate grant from the University of California–Santa Barbara. Robin Nagy is thanked for useful discussions and for providing fossil images. We thank Darlene Koerner and colleagues at Ashley National Forest and the Bureau of Land Management crew at John Jarvie Historic Ranch for their scientific and logistical support. The University of New Mexico provided C-isotope lab analyses. Karl Karlstrom, Laura Crossey, John Bloch, Viorel Atudorei, and Arlo Weil provided energetic field assistance. Bart Kowallis, Laura De Grey, and Andy Brehm provided photographs and mileage. REFERENCES CITED Ball, T.T., and Farmer, G.L., 1998, Infilling history of a Neoproterozoic intracratonic basin: Nd isotope provenance studies of the Uinta Mountain Group, western United States: Precambrian Research, v. 87, p. 1–18, doi: 10.1016/S0301-9268(97)00051-X. Bond, G.C., and Kominz, M.A., 1984, Construction of tectonic subsidence curves for the early Paleozoic miogeocline, southern Canadian Rocky

Mountains—Implications for subsidence mechanisms, age of breakup, and crustal thinning: Geological Society of America Bulletin, v. 95, p. 155–173, doi: 10.1130/0016-7606(1984)95<155:COTSCF>2.0.CO;2. Bryant, B., 1990, Geologic map of the Salt Lake City 30’ × 60’ quadrangle, north-central Utah, and Uinta County, Wyoming, with a section on palynologic data from Cretaceous and lower Tertiary rocks in the Salt Lake City 30’ × 60’ quadrangle: U.S. Geological Miscellaneous Investigations Series Map I-1944, 2 plates, scale 1:100,000. Bryant, B., 1992, Geologic and structure maps of the Salt Lake City 1° × 2° quadrangle, Utah and Wyoming: U.S. Geological Survey Miscellaneous Investigations Series Map I-1997, 2 plates, scale 1:250,000. Butterfield, N.J., 2000, Bangiomorpha pubescens n. gen., n. sp.: Implications for the evolution of sex, multicellularity and the Mesoproterozoic/ Neoproterozoic radiation of eukaryotes: Paleobiology, v. 26, p. 386–404. Butterfield, N., and Chandler, F., 1992, Paleoenvironmental distribution of Proterozoic microfossils, with an example from the Agu Bay Formation, Baffin Island: Paleontology, v. 35, part 4, p. 943–957. Colpron, M., Logan, J.M., and Mortensen, J.K., 2002, U-Pb zircon age constraint for the late Neoproterozoic rifting and initiation of the lower Paleozoic passive margin of western Laurentia: Canadian Journal of Earth Science, v. 39, p. 133–143. Condie, K.C., Lee, D., and Farmer, G.L., 2001, Tectonic setting and provenance of the Neoproterozoic Uinta Mountain and Big Cottonwood groups, northern Utah: constraints from geochemistry, Nd isotopes, and detrital modes: Sedimentary Geology, v. 141–142, p. 443–464, doi: 10.1016/S0037-0738(01)00086-0. Connor, J.J., Delaney, T.A., Kulik, D.M., Sawatzky, D.L., Whipple, J.W., and Ryan, G.S., 1988, Mineral resources of the Diamond Breaks Wilderness study area, Moffat County, Colorado, and Daggett County, Utah: United States Geological Survey Bulletin 1714-B, 15 p. Corsetti, F.A., and Kaufman, A.J., 2003, Stratigraphic investigations of carbon isotope anomalies and Neoproterozoic ice ages in Death Valley, California: Geological Society of America Bulletin, v. 115, p. 916–932. Crittenden, M.D., Jr., and Peterman, Z.E., 1975, Provisional Rb/Sr age of the Precambrian Uinta Mountain Group, northeastern Utah: Utah Geology, v. 2, no. 1, p. 75–77. De Grey, L.D., 2005, Geology of the Swallow Canyon 7.5-minute quadrangle, Daggett County, Utah and Moffat County, Colorado—Facies analysis and stratigraphy of the Neoproterozoic eastern Uinta Mountain Group [M.S. thesis]: Pocatello, Idaho State University, 122 p. De Grey, L.D., and Dehler, C.M., 2005, Stratigraphy and facies analysis of the eastern Uinta Mountain Group, Utah-Colorado border region, in Dehler, C.M., Pederson, J.L., Sprinkel, D.A., and Kowallis, B.J., eds., Uinta Mountain Geology: Utah Geological Association Publication 33, p. 17–33. Dehler, C.M., 2001, Facies analysis, cyclostratigraphy, and carbon-isotope stratigraphy of the Neoproterozoic Chuar Group, eastern Grand Canyon, Arizona [Ph.D. thesis]: Albuquerque, University of New Mexico, 371 p. Dehler, C.M., and Sprinkel, D.A., 2005, Revised stratigraphy and correlation of the Neoproterozoic Uinta Mountain Group, in Dehler, C.M., Pederson, J.L., Sprinkel, D.A., and Kowallis, B.J., eds., Uinta Mountain Geology: Utah Geological Association Publication 33, p. 35–48. Dehler, C.M., Elrick, M.E., Karlstrom, K.E., Smith, G.A., Crossey, L.J., and Timmons, J.M., 2001a, Neoproterozoic Chuar Group (~800–742 Ma), Grand Canyon: A record of cyclic marine deposition during global cooling and supercontinent rifting: Sedimentary Geology, v. 141–142, p. 465–499, doi: 10.1016/S0037-0738(01)00087-2. Dehler, C.M., Prave, A.R., Crossey, L.I., Karlstrom, K.E., Atudorei, V., and Porter, S.M., 2001b, Linking mid-Neoproterozoic successions in the western U.S.: The Chuar Group–Uinta Mountain Group–Pahrump Group connection (ChUMP): Geological Society of America Abstracts with Programs, v. 33, no. 5, p. 20–21. Dehler, C.M., Elrick, M.E., Bloch, J.D., Karlstrom, K.E., Crossey, L.J., and DesMarais, D., 2005, High-resolution δ13C stratigraphy of the Chuar Group (ca. 770–742 Ma), Grand Canyon: Implications for mid-Neoproterozoic climate change: Geological Society of America Bulletin, v. 117, no. 1/2, p. 32–45, doi: 10.1130/B25471.1. Dehler, C.M., Porter, S., De Gray, L.D., and Sprinkel, D.A., 2006, The Neoproterozoic Uinta Mountain Group revisited: A synthesis of recent work on the Red Pine Shale and related undivided clastic strata, northeastern Utah, in Link, P.K., and Lewis, R., eds., Proterozoic basins of Northwestern US: Society for Sedimentary Geology Special Publication (in press). Ehlers, T., and Chan, M., 1999, Tidal sedimentology and estuarine deposition of the Proterozoic Big Cottonwood Formation, Utah: Journal of Sedimentary Research, v. 69, no. 6, p. 1169–1180.

Neoproterozoic Uinta Mountain Group of northeastern Utah Fanning, M.C., and Dehler, C.M., 2005, Constraining depositional ages for Neoproterozoic siliciclastic sequences through detrital zircon ages: A ca. 770 Ma maximum age for the lower Uinta Mountain Group [abs.]: Geological Society of America Abstracts with Programs, v. 37, no. 6 (in press). Fanning, C.M., and Link, P.K., 2004, U-Pb SHRIMP ages of Neoproterozoic (Sturtian) glaciogenic Pocatello Formation, southeastern Idaho: Geology, v. 32, p. 881–884, doi: 10.1130/G20609.1. Hansen, W.R., 1965, Geology of the Flaming Gorge area Utah−Colorado−Wyoming: U.S. Geological Survey Professional Paper 490, 196 p. Hansen, W.R., and Rowley, P.D., 1991, Geologic map of the Hoy Mountain quadrangle, Daggett and Uintah Counties, Utah, and Moffat County, Colorado, United States: U.S. Geological Survey Map GQ-1695, scale 1:24,000. Hofmann, H.J., 1977, The problematic fossil Chuaria from the late Precambrian Uinta Mountain Group, Utah: Precambrian Research, v. 4, p. 1–11, doi: 10.1016/0301-9268(77)90009-2. Karlstrom, K.E., Bowring, S.A., Dehler, C.M., Knoll, A.H., Porter, S.M., DesMarais, D.J., Weil, A.B., Sharp, Z.D., Geissman, J.W., Elrick, M.B., Timmons, J.M., Crossey, L.J., and Davidek, K.L., 2000, Chuar Group of the Grand Canyon: Record of breakup of Rodinia, associated change in the global carbon cycle, and ecosystem expansion by 740 Ma: Geology, v. 28, p. 619– 622, doi: 10.1130/0091-7613(2000)028<0619:CGOTGC>2.3.CO;2. Kaufman, A.J., Knoll, A.H., and Narbonne, G.M., 1997, Isotopes, ice ages, and terminal Proterozoic earth history: Proceedings of the National Academy of Sciences of the United States of America, v. 94, p. 6600–6605, doi: 10.1073/pnas.94.13.6600. Knoll, A., Blick, N., and Awramik, S., 1981, Stratigraphic and ecologic implications of late Precambrian microfossils from Utah: American Journal of Science, v. 281, p. 247–263. Link, P.K., Christie-Blick, N., Devlin, W.J., Elston, D.P., Horodyski, R.J., Levy, M., Miller, J.M.G., Pearson, R.C., Prave, A., Stewart, J.H., Winston, D., Wright, L.A., and Wrucke, C.T., 1993, Middle and Late Proterozoic stratified rocks of the western U.S. Cordillera, Colorado Plateau, and Basin and Range province, in Reed, J.C., et al., eds., Precambrian; Conterminous U.S.: Boulder, Colorado, Geological Society of America, Geology of North America, v. C-2, p. 463–595. North American Stratigraphic Commission on Nomenclature, 1983, North American Stratigraphic Code: American Association of Petroleum Geologists Bulletin, v. 67, no. 5, p. 841–875. Nagy, R.M., and Porter, S.M., 2005, Paleontology of the Neoproterozoic Uinta Mountain Group, in Dehler, C.M., Pederson, J.L., Sprinkel, D.A., and Kowallis, B.J., eds., Uinta Mountain geology: Utah Geological Publication 33, p. 49–61. Nyberg, A.V., 1982a, Contributions of micropaleontology; Proterozoic stromatolitic chert and shale-facies microfossil assemblages from the western United States and the Soviet Union; morphology and relationships of the Cretaceous foraminifer Colomia Cushman and Bermudez [Ph.D. thesis]: Los Angeles, University of California, 265 p. Nyberg, A.V., 1982b, Proterozoic microfossils from the Uinta Mountain Group and from the Big Cottonwood and Mescal Formations of western North America: Geological Society of America Abstracts with Programs, v. 14, p. 220–221. Nyberg, A.V., Schopf, J.W., and Strathearn, G., 1980, Newly discovered shale facies microfossils from the Proterozoic Uinta Mountain Group, northeastern Utah: Geological Society of America Abstracts with Programs, v. 12, p. 299. Porter, S.M., 2004, The fossil record of early eukaryote diversification: Paleontological Society Papers, v. 10, p. 35–50. Porter, S.M., and Knoll, A.H., 2000, Neoproterozoic testate amoebae—evidence from vase-shaped microfossils in the Chuar Group: Grand Canyon: Paleobiology, v. 26, no. 3, p. 360–385. Porter, S.M., Meisterfeld, R., and Knoll, A.H., 2003, Vase-shaped microfossils from the Neoproterozoic Chuar Group, Grand Canyon—a classification guided by modern testate amoebae: Journal of Paleontology, v. 77, no. 3, p. 409–429. Prave, A.R., 1999, Two diamictites, two cap carbonates, two δ13C excursions, two rifts: The Neoproterozoic Kingston Peak Formation, Death Valley, California: Geology, v. 27, p. 339–342, doi: 10.1130/0091-7613(1999)027<0339: TDTCCT>2.3.CO;2. Sanderson, I.D., 1978, Sedimentology and paleoenvironments of the Mount Watson Formation, Upper Precambrian Uinta Mountain Group, Utah [Ph.D. dissertation]: Boulder, University of Colorado, 150 p.

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Sanderson, I.D., 1984, The Mount Watson Formation, an interpreted braidedfluvial deposit in the Uinta Mountain Group (upper Precambrian), Utah: The Mountain Geologist, v. 21, no. 4, p. 157–164. Sanderson, I.D., and Wiley, M.T., 1986, The Jesse Ewing Canyon Formation, an interpreted alluvial fan deposit in the basal Uinta Mountain Group (Middle Proterozoic), Utah: The Mountain Geologist, v. 23, no. 3, p. 77–89. Schell, E.M., 1969, Summary of the geology of the Sheep Creek Canyon Geological Area and vicinity, Daggett County, Utah, in Lindsay, J.B., ed., Geologic guidebook of the Uinta Mountains—Utah’s maverick range: Intermountain Association of Geologists and Utah Geological Society 16th Annual Field Conference, p. 143–152. Sears, J.W., Graff, P.J., and Holden, G.S., 1982, Tectonic evolution of lower Proterozoic rocks, Uinta Mountains, Utah and Colorado: Geological Society of America Bulletin, v. 93, p. 990–997, doi: 10.1130/00167606(1982)93<990:TEOLPR>2.0.CO;2. Sprinkel, D.A., 2002, Progress report geologic map of the Dutch John 30′ × 60′ quadrangle, Utah-Colorado-Wyoming (year 3 of 3): Utah Geological Survey Open-File Report 399, scale 1:62,500. Sprinkel, D.A., 2003, Geology of Flaming Gorge National Recreation Area, Utah-Wyoming, in Sprinkel, D.A., Chidsey, T.C., Jr., and Anderson, P.B., eds., Geology of Utah’s Parks and Monuments: Utah Geological Association Publication 28, 2nd edition, p. 277–299. Sprinkel, D.A., and Waanders, G., 2005, Organic microfossils and thermal maturation of the Neoproterozoic Uinta Mountain Group in the eastern Uinta Mountains, northeastern Utah, in Dehler, C.M., Pederson, J.L., Sprinkel, D.A., and Kowallis, B.J., eds., Uinta Mountain geology: Utah Geological Publication 33, p. 63–73. Sprinkel, D.A., Waanders, G., and Robbins, E.I., 2002, Stratigraphy, palynology, and maturity of the Proterozoic Uinta Mountain Group, eastern Uinta Mountains, Utah—Implications for unit thickness: Geological Society of America Abstracts with Programs, v. 33, no. 4, p. A-18. Sprinkel, D.A., Park, B., and Stevens, M.D., 2003, Geology of Sheep Creek Canyon Geological Area, northeastern Utah, in Sprinkel, D.A., Chidsey, T.C., Jr., and Anderson, P.B., eds., Geology of Utah’s Parks and Monuments: Utah Geological Association Publication 28, 2nd edition, p. 517–528. Stone, D.S., 1993, Tectonic evolution of the Uinta Mountains—palinspastic restoration of a structural cross section along longitude 109°15′, Utah: Utah Geological Survey Miscellaneous Publication 93-8, 19 p. Timmons, M.J., Karlstrom, K.E., Dehler, C.M., Geissman, J.W., and Heizler, M.T., 2001, Proterozoic multistage (ca. 1.1 and ca. 0.8 Ga) extension in the Grand Canyon Supergroup and establishment of northwest and north-south tectonic grains in the southwestern United States: Geological Society of America Bulletin, v. 113, p. 163–181, doi: 10.1130/00167606(2001)113<0163:PMCAGE>2.0.CO;2. Vidal, G., and Ford, T., 1985, Microbiotas from the late Proterozoic Chuar Group (northern Arizona) and Uinta Mountain Group (Utah) and their chronostratigraphic implications: Precambrian Research, v. 28, p. 349– 389, doi: 10.1016/0301-9268(85)90038-5. Vidal, G., and Nystuen, J., 1990, Micropaleontology, depositional environment, and biostratigraphy of the Upper Proterozoic Hedmark Group, Southern Norway: American Journal of Science, v. 290-A, p. 170–211. Wallace, C.A., 1972, A basin analysis of the upper Precambrian Uinta Mountain Group, Utah: Santa Barbara [Ph.D. dissertation]: Santa Barbara, University of California, 412 p. Wallace, C.A., and Crittenden, M.D., 1969, The stratigraphy, depositional environment and correlation of the Precambrian Uinta Mountain Group, western Uinta Mountains, Utah, in Lindsey, J.B., ed., Geologic guidebook of the Uinta Mountains: Intermountain Association of Geologists 16th Annual Field Conference, p. 127–142. Weil, A., Geissman, J., and Ashby, J.M., 2005, A new paleomagnetic pole for the Neoproterozoic Uinta Mountain supergroup, Rocky Mountain States, USA, Precambrian Research (in press). Williams, M.L., Crossey, L.J., Jercinoovic, M.J., Bloch, J.D., Karlstrom, K.E., Dehler, C.M., Heizler, M.T., Bowring, S.A., and Goncalves, P., 2003, Dating sedimentary sequences: In situ U/Th-Pb microprobe dating of early diagenetic monazite and Ar-Ar dating of marcasite nodules: case study from Neoproterozoic black shales in the southwestern U.S.: Geological Society of America Abstracts with Programs, v. 35, no. 6, p. 595. Williams, N.C., 1953, Late Pre-Cambrian and Early Paleozoic Geology of Western Uinta Mountains, Utah: American Association of Petroleum Geologists Bulletin, v. 37, no. 12, p. 2734–2742.

Printed in the USA

Geological Society of America Field Guide 6 2005

Basaltic volcanism of the central and western Snake River Plain: A guide to field relations between Twin Falls and Mountain Home, Idaho John W. Shervais Department of Geology, Utah State University, Logan, Utah 84322-4505, USA John D. Kauffman Idaho Geological Survey, University of Idaho, Moscow, Idaho 83844-3014, USA Virginia S. Gillerman Idaho Geological Survey, Boise State University, Boise, Idaho 83725-1535, USA Kurt L. Othberg Idaho Geological Survey, University of Idaho, Moscow, Idaho 83844-3014, USA Scott K. Vetter Department of Geology, Centenary College, Shreveport, Louisiana 71134, USA V. Ruth Hobson Meghan Zarnetske Department of Geology, Utah State University, Logan, Utah 84322-4505, USA Matthew F. Cooke Scott H. Matthews Department of Geological Sciences, University of South Carolina, Columbia, South Carolina 29208, USA Barry B. Hanan Department of Geological Sciences, San Diego State University, San Diego, California 92182-1020, USA

ABSTRACT Basaltic volcanism in the Snake River Plain of southern Idaho has long been associated with the concept of a mantle plume that was overridden by North America during the Neogene and now resides beneath the Yellowstone plateau. This concept is consistent with the time-transgressive nature of rhyolite volcanism in the plain, but the history of basaltic volcanism is more complex. In the eastern Snake River Plain, basalts erupted after the end of major silicic volcanism. The basalts typically erupt from small shield volcanoes that cover up to 680 km2 and may form elongate flows that extend 50–60 km from the central vent. The shields coalesce to form extensive plains of basalt that mantle the entire width of the plain, with the thickest accumulations of basalt forming an axial Shervais, J.W., Kauffman, J.D., Gillerman, V.S., Othberg, K.L., Vetter, S.K., Hobson, V.R., Zarnetske, M., Cooke, M.F., Matthews, S.H., and Hanan, B.B., 2005, Basaltic volcanism of the central and western Snake River Plain: A guide to field relations between Twin Falls and Mountain Home, Idaho, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 27–52, doi: 10.1130/2005.fld006(02). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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high along the length of the plain. In contrast, basaltic volcanism in the western Snake River Plain formed in two episodes: the first (ca. 7–9 Ma) immediately following the eruption of rhyolites lavas now exposed along the margins of the plain, and the second forming in the Pleistocene (≤2 Ma), long after active volcanism ceased in the adjacent eastern Snake River Plain. Pleistocene basalts of the western Snake River Plain are intercalated with, or overlie, lacustrine sediments of Pliocene-Pleistocene Lake Idaho, which filled the western Snake River Plain graben after the end of the first episode of basaltic volcanism. The contrast in occurrence and chemistry of basalt in the eastern and western plains suggest the interpretation of volcanism in the Snake River Plain is more nuanced than simple models proposed to date. Keywords: basaltic volcanism, basalt geochemistry, Snake River Plain, Bonneville flood.

INTRODUCTION The Snake River Plain of southern Idaho is one of the most distinctive physiographic features of North America (Fig. 1). The topographic low that defines the plain cuts across the structural grain of both the Idaho batholith and the Basin and Range province—even though formation of the plain coincided in time with Basin and Range extension. Evolution of the eastern Snake River Plain has been associated with the passage of North America over a mantle hotspot, forming a time-transgressive volcanic province that youngs from WSW to ENE (e.g., Morgan, 1972; Suppe et al., 1975; Armstrong et al., 1975). The onset of hotspot-related volcanism is marked by a series of overlapping caldera complexes, ignimbrites, and caldera-filling rhyolite lavas (Bonnichsen, 1982a, 1982b; Bonnichsen and Kauffman, 1987; Pierce and Morgan, 1992; Christiansen et al., 2002). The early rhyolite complexes were followed by extensive eruptions of plains basalts, which form a carapace on top of the earlier rhyolites (Malde and Powers, 1962; Greeley, 1982; Leeman, 1982; Kuntz et al., 1982,

45º 30’

W es

rn

P 2

SR

P

SR

MH

ste

n

3

Ea

te r

1 TF 41º 59’

42º 00’ 117º 02’

111º 05’

Figure 1. Physiographic map of the northwestern United States showing the Snake River Plain (SRP) and related features. Note the strong topographic expression of the plume track and the absence of Basin and Range extension across the axis of the plain. MH—Mountain Home; TF—Twin Falls.

1992; Malde, 1991). The hotspot is currently located under the Yellowstone Plateau, which also forms the locus of a gigantic geoid anomaly that underlies much of western North America (Smith and Braile, 1994; Pierce et al., 2002). The hotspot or plume model for the Snake River Plain is supported by studies of tectonic uplift and collapse along the plume track (Pierce and Morgan, 1992; Anders and Sleep, 1992; Smith and Braile, 1994; Rodgers et al., 2002), the 1000-km-wide geoid anomaly centered under Yellowstone (Smith and Braile, 1994), seismic tomography of the underlying mantle (Saltzer and Humphreys, 1997; Jordan et al., 2004), and helium isotopes (Craig, 1997). Alternate models have been proposed, however, such as localized asthenospheric upwelling associated with edge effects of North American plate motion, and counter flow created by the descending Farallon slab (Humphreys et al., 2000; Christiansen et al., 2002). There are two aspects of Snake River Plain volcanism that are not readily explained by the hotspot model. First, the eruption of basaltic lavas generally postdates passage of the hotspot in time, and may continue 2–3 m.y. after the onset of hotspot related volcanism farther to the NE. Second, the hotspot model does not explain the origin of volcanic rocks that do not lie on the presumed hotspot track; in particular, basalts of the western Snake River Plain cannot be directly related to the passage of North America over the Yellowstone hotspot, although models that link volcanism in the western Snake River Plain indirectly to the hotspot have been proposed (Geist and Richards, 1993; Camp, 1995; Shervais et al., 2002; Camp and Ross, 2004). Basalts of the Columbia Plateau, thought to represent the “head” of the Yellowstone plume, also lie well north of the presumed hotspot track. Several models have been proposed to explain this anomaly, including deflection of the plume by the Farallon plate (Geist and Richards, 1993), compression of the plume head by North American lithospheric mantle (Camp, 1995; Camp and Ross, 2004), or location of the plume head below northern Nevada (Pierce and Morgan, 1992; Pierce et al., 2002). The central Snake River Plain lies at the junction of the physiographic western and eastern Snake River Plain and represents a critical link between the two volcanic provinces. The

Basaltic volcanism of the central and western Snake River Plain structural, geophysical, sedimentary, and volcanic features of these provinces are distinct and require different origins. In this guidebook, we present a brief synopsis of geologic relations within and between the central and western Snake River Plain and then examine the stratigraphy and volcanology of each province in the field guide section. A compendium of recent papers on the Snake River Plain has recently been published by the Idaho Geological Survey (Bonnichsen et al., 2002), and interested readers are referred to that volume for more detailed information. THE SNAKE RIVER VOLCANIC PROVINCE: GEOLOGIC SETTING The central and western Snake River Plain comprise two distinct provinces with different crustal structure, stratigraphy, and volcanic history. Both provinces are characterized by crust that is thicker (40–45 km) than crust in the adjacent Basin and Range province (≈33 km; Mabey, 1978, 1982; Iyer, 1984; Malde, 1991). The western Snake River Plain is also characterized by a positive gravity anomaly that trends parallel to the axis of the plain, and magnetic anomalies that parallel its southern margin (Mabey, 1982). In contrast, gravity and magnetic anomalies in the eastern Snake River Plain are subdued and define a NW-trending texture that is normal to the trend of the eastern plain, but parallels structural trends in the adjacent Basin and Range province (Mabey, 1982; Malde, 1991). These contrasts reflect fundamental differences in the underlying structure and stratigraphy of the two terranes: the western Snake River Plain is a true graben, whereas the eastern Snake River Plain is structurally downwarped with little or no faulting along its margins (McQuarrie and Rodgers, 1998; Wood and Clemens, 2002; Rodgers et al., 2002). Central and Eastern Snake River Plain The ENE-trending central and eastern Snake River Plain begins as a gentle structural depression on the Owyhee Plateau that deepens to the NE and merges with the physiographic Snake River Plain near Twin Falls. This structural depression continues to the NE until it merges with the Yellowstone Plateau (Rodgers et al., 2002). The central Snake River Plain is defined loosely as that part of the eastern Snake River Plain trend that lies between the Owyhee Plateau and the Great Rift, or approximately between 116°W and 114°W along the axis of the plain. The increase in elevation of the eastern Snake River Plain from SW to NE is thought to result from thermal buoyancy in the upper mantle under the hotspot (Dueker and Humphreys, 1990; Pierce and Morgan, 1992; Smith and Braile, 1994). The progressive decay of this thermal anomaly with time has resulted in tectonic collapse in the wake of the “deformation parabola” that emanates from the hotspot (Anders and Sleep, 1992; Pierce and Morgan, 1992; Smith and Braile, 1994). The elevation difference between the Owyhee Plateau and the eastern Snake River Plain probably results from differences in the underlying crust: the Owyhee Plateau is underlain by the southern extension of the

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Idaho batholith, whereas the eastern Snake River Plain transects older crust of the Basin and Range province (Malde, 1991). The eastern Snake River Plain is characterized by 1–2 km of basalt that overlies rhyolite and welded tuff (e.g., Leeman, 1982; Kuntz et al., 1988, 1992; Greeley, 1982). Scientific drill holes at the Idaho National Laboratory (INL) site show that the basaltic suite ranges from <100 m to over 1500 m thick, with rhyolite basement extending to depths in excess of 3000 m (Malde, 1991; Hughes et al., 1999, 2002). Sedimentary intercalations consisting of fluvial sands, lacustrine muds, windblown sand, and loess range from 2 m to ≈25 m thick. The rhyolite eruptive centers consist of overlapping caldera complexes and ignimbrites that represent the initial volcanic activity at any given location along the axis of the Snake River Plain, and are thought to mark the arrival of the hotspot (e.g., Bonnichsen, 1982a; Christiansen, 1982; Draper, 1991; Morgan, 1992). Rhyolite eruptive centers become younger from SW to NE: the Bruneau-Jarbidge eruptive center (12.5–11.3 Ma), the Twin Falls eruptive center (10.0–8.6 Ma), the Picabo eruptive center (10.2 Ma), the Heise eruptive center (6.7–4.3 Ma), and the Island Park–Yellowstone eruptive center (1.8–0.6 Ma). The oldest basalts in the central and eastern Snake River Plain are slightly younger than the eruptive centers they mantle; the youngest basalts erupted from a series of NW-trending volcanic rift zones during the Holocene (e.g., the Shoshone lava flow, Craters of the Moon, the Great Rift, Hells Half Acre, Wapi lava flow), with flows as young as 2000 yr B.P. reported (Kuntz et al., 1986). Geophysical studies of the eastern Snake River Plain have documented differences in both the crustal structure and lithosphere-asthenosphere boundary relative to both the adjacent Basin and Range province and the Archean Wyoming craton (Braile et al., 1982; Iyer, 1984; Dueker and Humphreys, 1990, 1993; Humphreys and Dueker, 1994; Peng and Humphreys, 1998; Saltzer and Humphreys, 1997; Humphreys et al., 2000). The most significant crustal feature observed in seismic profiles of the eastern Snake River Plain is a midcrustal “sill” ≈10 km thick and 90 km wide that underlies the entire width of the eastern Snake River Plain, with seismic velocities (≈6.5 km/s) intermediate between the mafic lower crust and the more felsic intermediate crust (Braile et al., 1982; Peng and Humphreys, 1998). This mafic “sill” is inferred to represent basaltic melts that were intruded into the crust and either fractionated to form the high temperature rhyolites of the eastern Snake River Plain (McCurry and Hackett, 1999) or induced partial melting of the crust to form these rhyolites (Leeman, 1982; Bonnichsen, 1982a, 1982b; Doe et al., 1982). The crust is also underlain by a low velocity layer that is thought to be partially molten (Peng and Humphreys, 1998). This low-velocity layer apparently thickens to the NE, toward Yellowstone, where it dominates the lower crustal section (Priestley and Orcutt, 1982). Teleseismic tomography across the eastern Snake River Plain and adjacent domains show that the subcontinental lithosphere directly beneath the plain has been eroded to form a deep channel parallel to North American plate motion (Dueker and

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Humphreys, 1990; Humphreys and Dueker, 1994; Saltzer and Humphreys, 1997; Humphreys et al., 2000). The channel consists of low velocity, partially molten asthenosphere buttressed by levees of high-velocity, melt-depleted mantle. The depth of the lithosphere-asthenosphere boundary beneath the plain is roughly constant parallel to its axis (Saltzer and Humphreys, 1997). Since the subcontinental mantle lithosphere is thicker toward the NE, the sublithospheric channel must be more deeply eroded into the lithosphere to the NE. Tomographic images of the Yellowstone plume show that it dips steeply to the north from Yellowstone and extends to a depth of 400–600 km (Montelli et al., 2003; Jordan et al., 2004; Dueker et al., 2004). This is consistent with recent models of deep-seated mantle plumes, which show that they may tilt significantly from vertical and follow complex flow lines within the mantle (e.g., King, 2004; Farnetani and Samuel, 2004). Western Snake River Plain In contrast to the eastern Snake River Plain, the western Snake River Plain is a NW-trending structural graben, ≈70 km wide and 300 km long, bounded by en echelon normal faults and filled with up to 2–3 km of late Miocene through Pliocene sediment (Wood and Clemens, 2002; Shervais et al., 2002). Structural relief, based on deep drill holes that intercept basement, ranges from 2900 m near Mountain Home to over 3500 m on the southwest side of the valley (Malde, 1991). Sedimentary deposits in the western Snake River Plain range in age from Miocene through Quaternary; these deposits are dominantly lacustrine, with subordinate fluviatile and phreatomagmatic deposits (Wood and Clemens, 2002; Godchaux and Bonnichsen, 2002; Shervais et al., 2002). Miocene sediments were deposited in small, interconnected lakes, precursors to the large Pliocene “Lake Idaho” (Kimmel, 1982; Smith et al., 1982; Jenks and Bonnichsen, 1989; Malde 1991; Godchaux et al., 1992; Jenks et al., 1993; Wood and Clemens, 2002). The western Snake River Plain is topographically lower than the eastern Snake River Plain, with elevations ranging from 670 m to 1100 m (Malde, 1991), but recent data suggest that an ancestral Snake River system flowed southwards during the Miocene, implying higher elevations to the NW (Smith and Stearley, 1999; Link and Fanning, 1999). Paleontological evidence similarly suggests southward drainage in the late Miocene (Repenning et al., 1995). Volcanic activity in the western Snake River Plain began with the eruption of the silicic volcanic rocks along the northern and southern margins of the basin between ca. 11.8 and 9.2 Ma (Clemens and Wood, 1993; Wood and Clemens, 2002). Major basaltic activity in the western Snake River Plain occurred in two time periods: 9–7 Ma, and <2.2 Ma (White et al., 2002; Bonnichsen and Godchaux, 2002). The earlier episode is represented by basalt flows and phreatomagmatic vents intercalated with late Miocene sediments and by the acoustic basement that underlies much of the western graben (Malde and Powers, 1962, 1972; Amini et al., 1984; Malde, 1991; Wood and Clemens, 2002;

White et al., 2002; Shervais et al., 2002). The older (7–9 Ma) lavas and late Miocene to Pliocene sediments comprise the Idaho Group of Malde and Powers (1962). Young volcanic activity (<2.2 Ma) in the western Snake River Plain consists of: (1) plateau forming eruptions of tholeiitic basalt that form the volcanic uplands north and south of the Snake River, (2) tholeiitic shield volcanoes that sit on top of these uplands, and (3) young shield and cinder cone vents of alkaline to transitional alkaline basalt (Shervais et al., 2002; White et al., 2002). The younger lavas comprise the Snake River Group of Malde and Powers (1962) and correlate with the more abundant young volcanic rocks of the eastern Snake River Plain (Leeman, 1982). Volcanic rocks of similar age and character are also found in the Boise River drainage 40 km north of Mountain Home (Howard and Shervais, 1973; Howard et al., 1982; Vetter and Shervais, 1992). The physiographic junction between the western and eastern portions of the plain is located west of Twin Falls, Idaho (Fig. 1). Structurally, boundary faults of the western Snake River Plain graben extend SE into the eastern Snake River Plain plume track as far as Hagerman, and linears extend as far as Buhl, near Twin Falls. The axial Bouguer gravity high of the western Snake River Plain also extends into the plume track, and is deflected eastward south of the Mount Bennett Hills (Mabey, 1976, 1982). The positive magnetic anomaly along the SW margin of the western Snake River Plain curves to the NE where it intersects the plume track (Mabey, 1982), possibly outlining the southern margin of the Bruneau-Jarbidge eruptive center (Bonnichsen, 1982a). The NW-trending gravity and magnetic anomalies, which characterize the eastern Snake River Plain, are absent in this area. BASALTIC VOLCANISM: FIELD RELATIONS AND GEOCHEMISTRY Central Snake River Plain Late Neogene basalts of the central Snake River Plain north of Twin Falls form large shield volcanoes (Fig. 2) clustered along the axis of the plain, which overlie rhyolite of the Twin Falls eruptive center (Bonnichsen and Godchaux, 2002). This area has been mapped in detail during the past decade by the Idaho Geological Survey, by graduate students from the University of South Carolina and Utah State University, and by undergraduates from Centenary College. The results of this work have been compiled by the Idaho Geological Survey and are being prepared for publication as the Twin Falls 30′ × 60′ quadrangle geological map. Readers are referred to the preliminary copy of that map for the names, and extent of volcanic units in the central Snake River Plain and their relation to sedimentary units. All of the hills seen on this part of the plain are constructive volcanic vents. The shield volcanoes of the oldest vents have subdued topography and radial drainage and a thick cover of loess with well-developed soils, and they are typically farmed almost all of the way to the vent summits (e.g., Flat Top Butte, Johnson Butte). Many of the older vents appear smaller than the largest

Basaltic volcanism of the central and western Snake River Plain young vents because their flanks have been partially buried by younger lava flows (e.g., Skeleton Butte, Bacon Butte, Lincoln Butte). The surfaces of flows from younger vents are characterized by extremely rugged, chaotic topography, with inflated flow fronts, collapsed flow interiors, ridges, and collapse pits (e.g., Owinza Butte, Rocky Butte, Notch Butte, Wilson Butte). These surfaces lack established drainages but thin loess is present in local depressions. The surfaces of flows from vents of intermediate age have nearly continuous loess mantles of variable thickness and well-developed soils, but lack well defined surface drainage and are rarely farmed (e.g., Crater Butte, Dietrich Butte). Where the Bonneville Flood flowed overland, loess and soil were commonly stripped from basalt surfaces yielding surface morphology that appears younger then it should, which complicates identification of basalt units based on surface characteristics. Shield sizes vary and are difficult to constrain with confidence because the onlapping of younger lava flows often conceals the true size of the older shields. Even so, the largest are commonly 15–20 km across the main shield edifice—not counting elongate flows that may extend 50–60 km from the vent. Perhaps the most extensive shield in the region is Black Ridge Crater, located at the eastern edge of the central plain (113.96°W), which is ≈32 km across and covers ≈680 km2—all or parts of eight 7.5′ quadrangles. Shield volcanoes with extremely long flows include Black Ridge Crater, Notch Butte, Wilson Butte, and Rocky Butte. The Black Ridge Crater vent has a major sinuous lobe that flowed over 40 km from the vent, terminating just south of Notch Butte. Rocky Butte and Wilson Butte both fed extensive flow fields. Flows from Notch Butte extend west as far as Hagerman, a distance of ~42 km. A few radiometric dates are available for basalts in the central Snake River Plain (Armstrong et al., 1975; Tauxe et al., 2004; Idaho Geological Survey, 2005, unpublished data). These dates, along with stratigraphic relations, indicate that most of the shield volcanoes north of the Snake River are Quaternary, whereas those south of the river are early Quaternary or Tertiary. Older volcanoes north of the river, such as Lincoln Butte, Johnson Butte, and Sonnickson Butte, are surrounded and nearly buried by younger basalt flows, and typically have subdued morphology and radial drainage patterns. Remanent magnetic polarity of most of the Quaternary basalts north of the river is normal and within the Bruhnes polarity chron (<0.78 Ma); an exception is Sid Butte, which has reverse polarity, but because of its geomorphic characteristics it is thought to be early Quaternary and in the later part of the Matuyama chron. The older basalt units south of the river have both normal and reverse polarity and represent Tertiary to early Quaternary polarity events. Basalts of the central Snake River Plain have MgO similar to mid-ocean ridge basalt (MORB; 5%–10%) but with higher FeO* (12%–15%), TiO2 (1.6%–4.3%), P2O5 (0.4%–1.2%), and Al2O3 (14%–17%). They are also higher in FeO* than similar basalts of the eastern Snake River Plain (e.g., Idaho National Engineering Laboratory or INEL site) but less Fe-rich than typical ferrobasalts

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of the western Snake River Plain near Mountain Home (Fig. 3). The wide range in K, P, and K/P ratios at constant MgO implies a range of parent magmas derived from a similar source by variable degrees of partial melting. Fe8 values (≈13) imply deep melting or a source higher in FeO than MORB asthenosphere, while Na8 values (2.4–3.2) imply moderate but variable degrees of partial melting. Partial melting models based on 18 incompatible trace elements indicate 5%–10% melting of a spinel lherzolite source similar in composition to the E-MORB source. Garnet-bearing sources are ruled out by the slope of the rare earth element patterns, implying pressures less than 20–25 Kb, i.e., within the sublithospheric channel that has been imaged seismically. Most of the chemical variation within flows from single vents can be explained by low-pressure fractionation of the observed phenocrysts (olivine + plagioclase). Line scans of olivine phenocrysts show no reversals in composition or other evidence of magma mixing. Line scans of plagioclase phenocrysts show minor reversals that indicate fluctuations in magma chamber vapor pressures (Cooke, 1999; Matthews, 2000). The occurrence of cumulate gabbro xenoliths (ol+cpx+plg+oxide) in the Sid Butte vent is consistent with high-pressure fractionation at midcrustal levels, within the “basaltic sill” imaged seismically beneath the eastern plain (Matthews, 2000). The incompatible element ratios K/P and K/Ti decrease with mg# (= molar 100*Mg/[Mg+Fe]), ruling out extensive assimilation of older,

A

B

Figure 2. Field photographs contrasting the stratigraphy and topography of the central and western Snake River Plain. (A) Shield volcanoes of the central Snake River Plain north of Twin Falls. (B) Lacustrine sediments of the Glenns Ferry Formation near Mountain Home, overlain by plateau-forming Pleistocene basalts.

J.W. Shervais et al.

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Shoshone Mtn Home Tholeiites Mtn Home Alkali

4.5 4.0 3.5 3.0 2.5 2.0 17 16 15 14 13 12 17 16 15 14 13 12 11 12 11 10 9 8

3.5 3.0 2.5

1.0 .5

1.5 1.0 .5

350 300 250

60 50 40 30 400 350 300 250 200 150 10

9

8

7

6

5

MgO Figure 3. MgO variation diagrams for basalts of the central and western Snake River Plain; data from Cooke (1999), Matthews (2000), and unpublished results (Shervais). The tholeiitic basalts show extensive Fe and Ti enrichment not seen in the alkali series basalts, which are much higher in K. Variations in alumina at low MgO result from plagioclase accumulation or removal.

Basaltic volcanism of the central and western Snake River Plain felsic crust; these trends may be due to assimilation of previously injected gabbroic dikes at midcrustal depths. We infer that these basalts represent a mixed asthenospheric-lithospheric source that formed in response to the Yellowstone melting anomaly; these melts evolved by a combination of high-pressure and low-pressure crystal fractionation, with possible assimilation of previously intruded midcrustal ferrogabbros (Shervais et al., 2004). Western Snake River Plain The western Snake River Plain graben formed over a short time period ca. 10–12 Ma, coincident with the eruption of the extensive rhyolites that now form the flanks of this structure (Wood and Clemens, 2002). Late Miocene basalts (9–7 Ma) underlie the central part of the graben, as shown by geophysical surveys and deep drilling results (Wood, 1994; Wood and Clemens, 2002; Shervais et al., 2002). Overlying these basalts (and to some extent interbedded with them) are up to 2–3 km of lacustrine and fluviatile sediments that form the “Idaho Group” of Malde and Powers (1962): Chalk Hills Formation (oldest), Poison Creek Formation, Glenns Ferry Formation, and the Bruneau Formation. Phreatomagmatic vents intercalated with sediments of the Idaho Group are common in the area west of Mountain Home, between Grand View and Marsing along the Snake River (Godchaux et al., 1992; Godchaux and Bonnichsen, 2002; Bonnichsen and Godchaux, 2002). Mountain Home occupies an upland plateau above the Snake River floodplain, which is incised into fine-grained lacustrine sediments deposited by Lake Idaho during the late Miocene–early Pliocene (Fig. 2). Late Pliocene to early Pleistocene basalts sit conformably on the lacustrine strata, or occur within the uppermost few hundred feet. Flow of these basalts into the shallow margin of the lake resulted in deltas of hyaloclastite breccia and pillow lava, which pass upward into subaerial pahoehoe flows. Farther inland, flows are separated by deposits of fluvial gravel and sand. These plateau-forming basalts are overlain by a series of younger lavas erupted from 13 central vents that cluster near the NE margin of the plateau. The shield volcanoes rise 120–210 m above the surrounding plateau; several are capped by central depressions that probably represent former lava lakes (pit craters). Pleistocene cinder cones are the youngest volcanic features, and may cap small shield volcanoes. Recent tectonic activity is demonstrated by fault scarps with ~2–9 m of throw that crosscut all of the volcanic features. There are two dominant fault sets: one trends N85W, the other trends N60W (parallel to the range front faults) and appears to truncate against the N85W set. All of the faults are vertical or high angle normal faults that generally dip steeply to the south and are downthrown on the south or SW side (a few are downthrown to the N, forming small grabens with adjacent faults). The young faulting may be related to Basin and Range extension that has refracted to more westerly orientation to exploit preexisting fault zones of the western Snake River Plain. This relationship suggests that post–Lake Idaho volcanism in the western Snake River Plain

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may be related to reactivated partial melting of plume-modified lithosphere, in response to the onset of Basin and Range extension (Wood and Clemens, 2002; Shervais et al., 2002). Basaltic lavas of the western Snake River Plain are generally distinct from lavas of the eastern Snake River Plain trend, and are commonly ferrobasalts with up to 17% FeO* (Fig. 3; Shervais et al., 2002; White et al., 2002). This is clearly seen in the BruneauJarbidge eruptive center, where basalts of the eastern Snake River Plain trend are overlain by younger ferrobasalts of Salmon Falls Butte, which represents the SE extent of western Snake River Plain magmatism (Vetter and Shervais, 1997). In general, the oldest (pre–Lake Idaho) basalts are the most Fe-rich; younger basalts include ca. 2.0–0.7 Ma Fe-Ti basalts (equivalent to the Boise River Group 1 of Vetter and Shervais, 1992 and group M2 of White et al., 2002) and the <0.7 Ma alkali basalts which are high in K2O and MgO (equivalent to the Boise River Group 2 of Vetter and Shervais, 1992 and group M3 of White et al., 2002). This same sequence of basalts is observed in the Melba area, 100 km west of Mountain Home (White et al., 2002) and along the Boise River in the Smith Prairie area, 40 km north of Mountain Home (Vetter and Shervais, 1992). HYDROGEOLOGY OF BASALTIC VOLCANISM IN THE CENTRAL SNAKE RIVER PLAIN Detailed mapping in the central Snake River Plain has revealed new details of groundwater recharge and flow, and allows the formulation of new models for the formation of high conductivity aquifers. Groundwater flow in the Snake River Plain is commonly thought to be controlled by relatively porous horizons between lava flows (Lindholm and Vaccaro, 1988; Welhan et al., 2002c). These so-called interflow zones are characterized by complex geometries reflecting the fractal nature of the pahoehoe inflationary process, the quasi-stratigraphic layering and interfingering of flows within and between lava fields (shield volcanoes), and the development of tension fissure networks along the margins of inflated flow structures (Welhan et al., 2002b). These permeable zones can be mapped in the subsurface using core and geophysical logs and their autocorrelation length scales used to construct stochastic models of subsurface pathways of preferential flow (Welhan et al., 2002a). Direct evidence of the “pipeline” nature of this flow, as reflected in extremely low dispersivities associated with mass transport on scales of 25 km, has been documented by Cecil et al. (2000) using chlorine-36. The mapping of Cooke (1999) and Matthews (2000) shows that thick alluvial intervals are most common where adjacent volcanic edifices abut one another and overlap. These overlapping flows create moats, which control the location of surface stream drainages and fill with coarse alluvium. Younger lava flows are channeled into these drainages, displacing the streams and covering the alluvium with relatively thick, semi-permeable caps, forming elongate aquifers with extremely high conductivities that follow the preexisting paleo-drainage. Reconstruction of

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basaltic volcanism through time using detailed geologic maps allows us to predict the location of paleo-drainage and elongate alluvial aquifers. This ability to predict aquifer location will prove to be a valuable tool as increased demands are placed upon the groundwater supply by agriculture and population growth. In addition, our mapping reveals that young basalt vents seldom exhibit channelized drainage systems that connect with major through-going streams. Instead, the rugged volcanic topography of ridges, flow fronts, lava channels, and collapse pits trap precipitation, which must either evaporate or percolate into the fractured lavas to recharge the local groundwater. These young volcanic features constitute “negative basins” of interior drainage, despite their topographic emergence. We suggest that these young basalts represent recharge zones that can be easily mapped and distinguished from older flows that display well-developed external drainage (Cooke and Shervais, 1999). On a more regional scale, the eastern and central Snake River Plain basalts and intercalated units form the Snake River Plain aquifer, which along with the Snake River system supplies water to southern Idaho and are essential to the agricultural development of the state. Water recharges from snow melt in the mountains north of the plain and in local basalt recharge areas. Groundwater flow is from the NE to SW and it discharges as springs in the Thousand Springs and other reaches of the north and east wall of the Snake River Canyon between Twin Falls and King Hill. The spring water is used extensively by the aquaculture industry, but quantity and quality of the water has been declining due to increased groundwater pumping upgradient, drought, and changes in surface irrigation systems. Spring and groundwater irrigation is supplemented by surface withdrawals from the Snake River, but acute water shortages are posing a political dilemma for Idaho politicians and water experts. Mapping of springs in the northwall of the Snake River Canyon by Covington and Weaver (1989), the Idaho Geological Survey, and others suggests a geologic control. Covington et al. (1985) (and others before them) recognized that some springs discharge at rubble zones at the base of where canyon-filling late Quaternary basalts entered a paleodrainage system. It also appears that the topography developed on the older, slightly altered (Banbury) basalts may have influenced spring discharge, as many springs seem to emerge through talus but just above the QuaternaryTertiary unconformity. The older units may locally form a more impermeable horizon and basal aquitard. FIELD TRIP GUIDE Day 1 (Half Day) Hagerman Valley Directions to Stop 1.1 From Hagerman, drive south on Idaho Highway 30 ~7 mi (11 km) across the bridge over the Snake River to ~1 mi south of Thousand Springs Resort. Pull off to the right on a small dirt road across from the entrance gate to a big house on the river. Stop 1 lies ~200 ft (60 m) uphill to the right. Bonneville flood effects in

the Hagerman Valley are shown on Figure 4; all Day 1 stops are shown on Figure 5. On the way to Stop 1, we will drive by scour features and deposits created by the Bonneville Flood. The following paragraphs briefly describe some of the effects of the floodwaters (Fig. 4). About fifty years ago, geologists first recognized that giant gravel bars and stark erosional features common along the Snake River were caused by cataclysmic lowering of Pleistocene Lake Bonneville. The most complete descriptions and analyses of the Bonneville Flood were written by Malde (1968) and O’Connor (1993). O’Connor (1993, p. 1) writes, “Approximately 14,500 years ago, Pleistocene Lake Bonneville discharged 4750 km3 of water over the divide between the closed Bonneville basin and the watershed of the Snake River. The resulting flood, released near Red Rock Pass, Idaho, followed the present courses of March Creek, the Portneuf River, and the Snake and Columbia River before reaching the Pacific Ocean. For 1100 km between Red Rock Pass and Lewiston, Idaho, the Bonneville Flood left a spectacular array of flood features.” O’Connor (1993) goes on to describe the hydrology, hydraulics, and geomorphology of the flood and provides a picture of the history and character of the flood and its many landforms and deposits. A major characteristic of the flood is a stepwise drop in the water surface caused by the diverse geologic and geomorphic environments along the flood route. A repeated phenomenon is a narrow canyon with constricted but high-discharge flow that opens up into a wider valley with less restricted, lower-discharge flow. In many locations, a constriction back-watered the flood causing inundation and overland flow. East of Twin Falls flood water diverted north of the Snake River canyon scoured a long stretch of basalt surfaces and rejoined flood waters in the canyon for ~17 km forming several systems of cataracts and plunge pools (see Day 3, Stop 3.6). The areas of Hagerman and Thousand Springs (Day 1) show the stepwise nature of the flood’s energy and the resulting features. The Thousand Springs reach was a constriction that overfilled and the overland flow scoured and plucked basalt at the Sand Springs Nature Preserve, Box Canyon, and Blind Canyon. South of Thousand Springs a dry valley appears to have been the previous location of the confluence of Salmon Falls Creek and the Snake River. The flood may have scoured a new channel for the Snake River and subsequently the river abandoned the previous valley. Just downstream of Thousand Springs, where the Hagerman valley opens up, energy dropped and the floodwaters deposited a large expansion bar 5 km long with 4 m boulders at its upstream end. At maximum discharge, the floodwaters would have been ~60 m deep at the location of the Hagerman Inn. In the bottom of the valley, thick deposits of Yahoo clay that had buried basalt of Notch Butte were stripped away and the surface of the basalt was scoured. Stop 1.1—Banbury Basalt and Bacon Butte Basalt: Basalt Water Interaction [N 42°43.804′, W 114°51.079′] Mapping and reexamination of the “older” volcanic units exposed south of Hagerman is one phase of the Idaho Geological

Basaltic volcanism of the central and western Snake River Plain

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Figure 4. Topographic map of Thousand Springs and the southern Hagerman Valley showing erosional and depositional features of the Bonneville Flood. Large blue arrows show inferred floodwater flow directions. The arrow pointing northwest is placed to show the change from scoured rock to the bouldery expansion bar deposited in the broadening valley.

Survey’s STATEMAP project in the Twin Falls 30′ × 60′ quadrangle. The area includes Banbury Hot Springs, type locality for the Tertiary Banbury Basalt, which Malde et al. (1963) mapped over a wide area of southwestern Idaho. Our work and the excellent mapping of Malde and Powers (1972) show much heterogeneity within the “type section.” The lower Banbury Basalt

includes a field of hydrovolcanoes, which appear to be localized along a NW-trending structural zone that also influenced later graben development and formation of Pliocene Lake Idaho, geothermal activity, and canyon development. The inferred structure is suspiciously close to a projection (S49E) of the northern margin of the western Snake River Plain.

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Stop 1.1 42º43’ 48” N, 114º 51’ 05” W Stop 1.2 42º43’ 35” N, 114º 51’ 21” W Figure 5. Topographic map of the Snake River canyon between Banbury Springs and Sand Springs showing locations for Day 1 of the field trip, and erosional and depositional features of the Bonneville Flood. Relict flood features include scoured and plucked basalt and abandoned valley with no established drainage.

Stop 1.3 42º43’ 11” N, 114º 51’ 24” W

Stop 1.4 42º42’ 24” N, 114º 50’ 23” W

Stratigraphically overlying Miocene rhyolite, the MiocenePliocene Banbury Basalt consists of a lower sequence of altered basalt flows and vent deposits, a middle layer of basaltic pyroclastic deposits overlain by and/or interbedded with sediments, which thicken to the south, and an upper sequence of basalt flows (Malde and Powers, 1972). The vent facies of the lower Banbury Basalt includes at least two exposed volcanic centers (Thousand Hill and Riverside Ferry vents), which are characterized by tuff breccias with laterally extensive and locally palagonitized surge and tephra deposits. At Stop 1.2, spectacular 10-m-high beds of tuff breccia contain blocks larger than 1 m. At Stop 1.1, distal phreatomagmatic tuffs containing volcanic bombs, stream pebbles, and siltstone xenoliths are overlain by a 1-m-thick bed of black spatter. The spatter is overlain by an oxidized tuff with small glass bombs, and above that are three upward-coarsening cycles of air-fall tuff. Unconformably overlying the emergent hydrovolcano is a baked siltstone and the upper Banbury Basalt, which here consists of notably fresher, plagioclase-phyric, vesicular basalt flows. Elsewhere in the map area, the upper Banbury consists of altered or water-affected basalt (WAB of Godchaux and Bonnichsen, 2002) with a different magnetic polarity than the feldspar-rich flows. Pending geochronology and chemistry may improve age constraints and correlations. Even at its type locality, the Banbury Basalt is lithologically heterogeneous with lateral facies changes indicating basaltic volcanism within a

lacustrine and fluvial setting prior to Lake Idaho and probably over a considerable time span (Gillerman, 2004). The heterogeneity, polarity reversals, and spatial distribution of the flows also indicate a variety of sources for the Banbury units. As noted by others, it is time to revise the nomenclature. The recent Idaho Geological Survey mapping, which will be displayed on the field trip, has renamed and more precisely subdivided many units based on mapping and paleomagnetism, combined with chemistry and a very few radiometric age dates. Stop 1.1A—Exposure of the Lower Banbury Distal Vent Facies Volcaniclastics The prominent outcrops to the right of the gully are massive basaltic tuffs in 3–4 upward-coarsening cycles, 2–3 m thick, with laminated to cross-bedded tuff above and thin (6–8 cm), locally cross-bedded, fine tuff below (Fig. 6A). This sequence overlies a juvenile conglomerate of basalt, hyaloclastites, and some rounded stream cobbles, that transitions up to a bleached/oxidized (?) pumice bed (Fig. 6B). Below these beds are vent-like spatter and pillowy basalt with abundant fresh glass. Although the base of the sequence here is unexposed, Malde and Powers (1972) mapped the lower portion of the slope as lower Banbury sediments and tuff, most likely from better exposures ~1 mi to the north, near the Sligar’s Thousand Springs Resort. There, a

Basaltic volcanism of the central and western Snake River Plain section of mixed tuffs, sandstones, and massive basaltic units are exposed in a gully and in old road cuts. The massive brown and black mottled basalts are coarse grained, magnetic, and look almost gabbroic. An exposure of the mottled rock is just below the Stop 1.1B outcrop. Godchaux and Bonnichsen (2002) referred to these massive mottled rocks as “spotted” due to black augite crystals in a lighter brownish groundmass, which might include glass and fibrous and hydrous minerals such as zeolites, chlorite, or amphiboles. They interpret the spotted, massive lavas as water-affected basalt, which formed in deep water (over a few tens of feet). If so, then the volcaniclastic sediments should be indicative of such an environment. It may be that volcanic eruption of flows, tectonic activity, and climatic cycles created a rapidly changing and laterally heterogeneous distribution of lakes and sedimentary facies during deposition of the “Banbury Basalt” and related units. At this stop, the spatter layers dip into the hill (strike N65E, dip 11°NW to EW and 14°N) and as seen from a distant photo (Fig. 6C) the entire lower Banbury Basalt section (renamed Tvd or vent deposits, distal), including the baked siltstone above, appears to be tilted 10–15°NNW, possibly faulted along a WNW normal fault. At this location, the upper Banbury lavas are clearly unconformable on top of the lower Banbury volcaniclastics. The upper Banbury lavas are flat lying and mostly subaerial, but the lower part is glassy, popcorn-like, with quasi-pillows and glass. The flows here are distinguished by abundant feldspar phenocrysts and normal magnetic polarity, and have been renamed the Basalt of Oster Lakes (Tob) for exposures farther east and several hundred feet topographically lower. Structural relationships, age, and exact correlations of these various exposures are problematic.

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A

B

C

Stop 1.1B—Quaternary Basalt of Bacon Butte: Water Escape Features The Basalt of Bacon Butte (formerly included in Sand Springs basalt by Malde and Powers, 1972) fills paleo-drainage here, with evidence for water interaction at base (pillows and hyaloclastites) and vertical dewatering conduits that penetrate entire flow (Fig. 7). The water-escape structures in this flow constitute a feature of volcanic rocks that have not been described previously. They consist of subvertical conduits ≈1 m in diameter (0.5–1.5 m range) and spaced 5–15 m apart, partly filled with spicaceous basaltic rubble. Columnar jointing in the adjacent basalt wraps from vertical along the base of the flow to horizontal adjacent to the conduits. Vesicles are more common in the wallrock of the conduits than in parts of the flow farther away, and the basalt is quenched to a glassy or aphanitic texture in the wallrock adjacent to the conduit (Fig. 7A). In many of the conduits, basalt from the adjacent wallrock forms a lattice or trellis arrangement within the conduit that is contiguous with basalt of the wallrock (Fig. 7B). The basalt in this trelliswork has glassy margins and spikey projections along its surface. Spicaceous rubble is common in all conduits, often forming discrete fragments that resemble

Figure 6. Banbury Basalt in Hagerman Valley. (A) Lower Banbury Basalt, volcaniclastics and phreatomagmatic vent facies (Stop 1.1). (B) Bleached pumice bed in lower Banbury Basalt. (C) Unconformity of upper Banbury Basalt (Basalt of Oster Lakes) over tilted section of lower Banbury phreatomagmatic volcaniclastics with primary and probable structural dip (Stop 1.1). Fence posts for scale.

popcorn. This spicaceous rubble of “popcorn” basalt does not represent later fill or pieces broken from the wallrock, but rather blobs of lava that were isolated within the conduit while steam was actively rising through the flow (Fig. 7C). The upper surface of the lava flow contains depressions as much as 10 m across and 1.5 m deep centered above the waterescape structures, which appear to represent areas where the surface lava collapsed into the conduit during water escape. Such

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Figure 7. Water-escape structures in basalt of Bacon Butte, Stop 1.1B. (A) Panorama of outcrop showing water-escape conduits, rotated columnar jointing, and rubbly fill. (B) Close-up of water escape conduit showing trellis structure and spiny rubble with popcorn texture. (C) Popcorn texture basalt in water conduit. (D) Large (1–2 m) scale steam cavities at base of flow, just above section of megapillows.

a feature is well illustrated at another location by the occurrence of a ropey pahoehoe flow surface that is preserved deep inside a lava flow within a steam conduit. The subaerial lava flow with water escape structures is underlain by pillow lava that fills intervening depressions in the paleosurface. Many of these are mega-pillows as much as 4 m across, which may contain water-escape structures that continue up into the subaerial flow. In addition, the interface between the pillow zone and the subaerial flow commonly contains small caves and cavities that represent large steam bubbles (>1 m across) trapped below the lava flow (Fig. 7D). Directions to Stop 1.2 Continue South 0.3 mi (0.5 km) on Hwy 30 to a prominent draw on the right with several dirt roads. Turn in to the right, go through gate and close it, and park near end of drivable jeep trail to right. Hike northward on cow trail along wire fence ~1/3 mi

(0.5 km) to farthest gully with trees and brush. Hike steeply uphill 150′–200′ (45–60 m) to large, cliffy exposure of the megabreccia at top of ridge on north side of gully. Stop 1.2—Banbury Basalt: Megabreccias [N 42°43.583′, W 114°51.35′] As you climb up to the ridgetop, there are poor exposures on the slope of white breccias and palagonite tuffs, as well as lavas. The outcrop on the ridge top is an ~10-m-thick “megabreccia” unit, which can be followed ~600′ to the west, where it terminates, perhaps on the crater rim, or is faulted. The megabreccia (Tvp) consists of unsorted tuff breccia with large accidental blocks, many more than 1 m in length (Fig. 8A). Block compositions include several types of basalt (lower Banbury Basalt in appearance), tan fine-grained lake sediment, sparse stream pebbles, and some juvenile spattery tephra.

Basaltic volcanism of the central and western Snake River Plain

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Figure 8. Vent facies of the Banbury Basalt. (A) Megabreccias that probably represent a vent facies (Stop 1.3). (B) Riverside Ferry vent (Stop 1.4).

Some clasts show evidence of hot reaction rims on spheroidally weathered basalt. The matrix is unstratified with ~25% fine ash and many small (<1cm) clasts. The breccia unit appears to grade upwards into tuff with fewer large clasts. The megabreccia and associated tuff are interpreted as explosive, subaerial crater-filling deposits of the Thousand Hill vent. This small drainage basin (~1/2 square mile) is characterized by outcrops of megabreccia, tuffs, red-orange sediment (which always overlies the volcaniclastics), and lavas, and by faults. The area was mapped by Stearns et al. (1938) as “Fault complex of Hagerman lake beds and interstratified tuff.” Malde and Powers (1972) mapped it as lower Banbury “vent deposits of basaltic pyroclastic material,” without showing any internal detail. The geology is very complex and diverse, and the volcanic and structural relationships are obscured by a veneer of tan-colored Quaternary Yahoo Clay (Bruneau sediments in earlier work), which unconformably overlies the Tertiary volcanics. Directions to Stop 1.3 Return toward vehicles and hike over to small hill nearby, labeled “Thousand” on topographic map. Stop 1.3—Banbury Basalt: Tilted Hydromagmatic Vent and Vent Facies Breccias [N 42°43.183′, W 114°51.40′] The outcrop consists of a 7-m-thick, unsorted and unstratified, matrix-supported volcanic breccia with dull olive-green to tan color containing accidental blocks of vesicular basalt (as large as 0.5 m across) and numerous polished stream pebbles. Note the oblong blocks of undisturbed, fine-grained tan water-laid sediment, which must have been horizontal originally, suggesting that dip is the result of structural tilting. This is a section of the vent deposits, dipping 40 degrees north, probably due to faulting. Hike around to top surface (now dipping) to look at strong linear features, which are interpreted as elutriation (gas escape) features. This outcrop may be the same unit as in Stop 1.2, but on the opposite side of vent. Note that the Thousand Hill vent and the Riverside Ferry vent (Stop 1.4) are both along a strong northwest-trending structural orientation.

Directions to Stop 1.4 Drive south on Hwy 30, ~2 mi (3.2 km), and turn left at intersection with paved “River Road.” Follow River Road past houses, ~1 mi (1.6 km) and turn left on paved road near power line. Drive past sequence of orange-colored palagonitic tuffs on right roadcut and continue to gravel road at top of hill. Turn right on gravel road, go ~1/3 mi (0.5 km) and fork left. Drive carefully northward to edge of vent exposure. Some roads are very sandy and may need 4WD. There are at least two jeep trails that access the vent area and overlook the Snake River. Stop 1.4—Banbury Basalt: Riverside Vent (optional, if time permits) [N 42°42.40′, W 114°50.383′] This exposure in the north half of the section 29 “island” is a hydrovolcano bisected by the Snake River and originally named the Riverside Ferry cone by Stearns et al. (1938). Bonnichsen and Godchaux (2002) referred to it as the Riverside Ferry cone but named the phreatomagmatic units “Tuff of Blue Heart Springs” after a spring located on the north side of the river adjacent to the volcano. They described the tuff as explosively erupted layers of cinder to spatter-rich deposits tilted near the vent constructs and transitioning to subhorizontal, finer-grained units more distally. Chemically they place the basaltic material in the general Snake River olivine tholeiite compositional range. They also place some of the Banbury tuff units as intercalating with the Glenns Ferry Formation sediments, although the mapping of Malde and Powers (1972) would generally place the Banbury Basalt as older than the thick pile of lake sediments. Certainly there are “Banbury-age” lake and fluvial sediments interstratified with basaltic units (4–6 Ma) and there are basaltic tuffs and basalt (3.4–3.8 Ma) interbedded within the Glenns Ferry Formation at the Hagerman Fossils National Monument (W.K. Hart and M.E. Brueseke, 1999, personal commun.). Dating the Banbury units is difficult due to the alteration, although a few recent Ar-Ar radiometric dates are available, but not definitive. The core of the Riverside Ferry vent is well exposed by erosion (Fig. 8B). A massive tuff breccia, air-fall tuffs and breccia, cindery beds, surge deposits, and palagonitized tuffs can be found within a short hike across the 1000 ft of exposed vent.

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On the southeast side of the vent, a basaltic lava flow appears to overlie the vent deposits and may be related. Exposures on the north side of the river contain abundant red cindery material, but these deposits have not been studied in detail. A NWtrending fault may underlie the course of the Snake River, and it is probable that there is a significant northwest-trending fault under the “dry” river course south of the “island.” Evidence for this structure(s) includes small visible offsets of the older rhyolites and lower Banbury lavas, topographic features emphasized by the Bonneville Flood (Fig. 4), and alignment of geothermal wells and hot springs. Return to Hagerman on Hwy 30. Day 2 (All Day) Mountain Home Directions to Stop 2.1 Drive north on I-84 to Mountain Home; take Exit 95 on Idaho Hwy 20 north toward Camas Prairie and Sun Valley. Continue N for ~7 mi (11 km) where the highway enters Rattlesnake Creek canyon; park at the turnout on the right side of highway. All Day 2 stops are shown on Figure 9A; detailed topography for stops 2.2 through 2.4 is shown in Figure 9B. Stop 2.1—Danskin Mountains Rhyolite, Highway 68 North of Mountain Home [N 43°12.052′, W 115°33.2′] The oldest volcanic rocks exposed in the Mountain Home area are rhyolite lava flows that form the Danskin Mountains and the Mount Bennett Hills. Clemens and Wood (1993) mapped the rhyolite here as the Danskin Mountains Rhyolite, and mapped rhyolite which is exposed farther east as the Mount Bennett Rhyolite. They determined a K-Ar age of 10.0 ± 0.3 Ma for sanidine in a vitrophyre from the summit area on Teapot Dome (Danskin Mountains Rhyolite), and a K-Ar age of 11.0 ± 0.5 Ma for plagioclase in a rhyolite from near Mount Bennett. Possible correlative units on the south side of the western Snake River graben include the Sheep Creek Rhyolite (9.88 ± 0.46 Ma; Hart and Aronson, 1983), the rhyolite of Tigert Springs, the rhyolite of O X Prong, and the rhyolite of Rattlesnake Creek (Kauffman and Bonnichsen, 1990; Jenks et al., 1993). Rhyolite in the eastern Mount Bennett Hills ranges in age from 9.2 ± 0.13 Ma to 10.1 ± 0.3 Ma (Armstrong et al., 1980; Honjo et al., 1986, 1992). The Danskin Mountains Rhyolite is typically a vitrophyre with abundant phenocrysts of sanidine and quartz set in a red, gray-brown, or black volcanic glass. Flow banding appears as laminar variations in the color of the glass, or in its crystallinity. The flow banding is commonly folded ptygmatically, indicating rheomorphic mobilization of the rhyolite. Flow banding and axial foliation in the vitrophyre generally trend N50ºW to N60ºW, and dip 15º–45°NE. There are no indications that these rhyolites are rheomorphic ignimbrites; they appear to be rhyolite lavas erupted from fissures that were subparallel to the current range-front faults (Bonnichsen, personal commun., 1996, 1997).

Directions to Stop 2.2 Return to I-84 and proceed west toward Boise; take the next available exit (exit number 90) then continue across and back on to I-84 toward Boise (this will slow you down and position you for the next stop). Pull off to the right immediately after entering the freeway at the first roadcut encountered. Stop 2.2—Plagioclase Flotation Cumulates of Lockman Butte, I-84 Roadcut [N 43°10.599′, W 115°45.601′] This roadcut transects a major flow lobe from Lockman Butte that consists of multiple lava flows that are stacked one on top of (or beneath) another. Many of these flow units contain plagioclase flotation cumulates that formed within covered lava channels and aphyric ferrobasalt residues that drained out the bottom of the cumulates (McGee and Shervais, 1997; Shervais et al., 2002; Zarnetske and Shervais, 2004). Another interesting aspect of these flows are the gigantic vesicles found in some gas accumulation horizons; these vesicles are up 2 m long and 0.6 m high—big enough to lie down in! Lava flows with flotation cumulates comprise three zones: a central, plagioclase porphyry with intersertal to intergranular textures in the groundmass, an upper diktytaxitic zone comprising plagioclase laths with large voids and minor intergranular mafics and glass, and a lower aphyric zone of ferrobasalt, with 16%–17% FeO* (Fig. 10A). Mass balance calculations show that the diktytaxic zone contains 30%–50% porosity, represented by the ferrobasalt base, if the central plagioclase porphyry is assumed to represent the bulk composition. Detailed outcrop maps show that successive lava flows flowed beneath previously emplaced flows, inflating and plastically deforming the aphyric ferrobasalt zone in the overlying flow. Plagioclase flotation cumulates beneath this ferrobasalt ceiling display a horizontal contact between the diktytaxitic and plagioclase porphyry zones; we interpret this horizon to represent the contact between interstitial melt (below) and volcanic gasses (above), and suggest that interstitial melt was displaced by the rising gasses (McGee and Shervais, 1997; Zarnetske and Shervais, 2004). Plagioclase phenocrysts are An65 in both the porphyry and diktytaxitic zones, suggesting formation of the plagioclase framework by simultaneous crystallization throughout the flow, followed by sinking of the dense, interstitial ferrobasaltic liquid to the bottom of the flow. This interpretation is supported by olivine phenocrysts (Fo68) trapped in the plagioclase framework, and by partially replaced intergrowths containing Fo80–89 olivine. The Fo-rich olivine relicts imply a parent magma that was significantly more magnesian than observed. These flotation cumulates provide an analogue for anorthosite formation, both in the lunar magma ocean and in Earth’s mid-crust. We suggest that planetary anorthosites represent plagioclase saturated melts that crystallize en mass to form buoyant rafts of plagioclase plus trapped mafics; interstitial ferrobasalts liquids are forced out by rising magmatic gasses, and sink back into the underlying magma (McGee and Shervais, 1997; Zarnetske and Shervais, 2004).

Basaltic volcanism of the central and western Snake River Plain

A

Stop 2.3

Stop 2.2 Stop 2.1

Stop 2.4

Stop 2.5

B

Stop 2.3 Stop 2.2 Stop 2.4

Figure 9. (A) Topographic map of the western Snake River Plain around Mountain Home, Idaho, showing Day 2 stops. (B) Detail of Stops 2.2 through 2.4 NW of Mountain Home.

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Directions to Stop 2.3 Drive NW on I-84 for ≈14 mi (23 km) to the first exit (Simco Road, Exit 74); cross the freeway and return to I-84 in the SE direction. Exit at the first Mountain Home exit (exit number 90) and take the first right turn encountered on Business I-84 (old Hwy 30). Proceed along the frontage road parallel to the railroad tracks for ≈6 mi (≈10 km), then turn left and cross the RR tracks at a clearly marked crossing. Proceed ≈1.8 mi (2.9 km) along this wide gravel road to a dirt road turn-off on the left near the crest of the grade. Turn left and follow this dirt road ≈1.3 mi (2.1 km) to the rim of Crater Rings. You will go past a stock pond (commonly dry) and through at least one wire fence; be sure to close all gates behind you. Note: Crater Rings is rattlesnake heaven, so watch where you step and put your hands!

A

Stop 2.3—Crater Rings Pit Craters: Lava Lakes and Spatter Ramparts [N 43°11.715′, W 115°50.866′] The Crater Rings are a spectacular volcanic feature that consists of two large pit craters at the summit of a broad shield volcano (Fig. 10B). The western crater is ≈800 m across and 75 m deep; the eastern crater is ≈900 m across and 105 m deep. The inner walls of the pit craters consist of welded spatter, agglutinate, and minor intercalated lava flows; no fragmental horizons are exposed in the crater walls, which argues against phreatomagmatic eruption. The welded spatter and agglutinate are easily identified by their characteristic textures, oxidized coloration, and hollow ring when struck with a hammer. The eastern crater is surrounded on three sides by spatter and agglutinate ramparts. The Crater Rings represent pit craters that were filled episodically with lava lakes (Shervais et al., 2002). They are equivalent to similar features in Hawaii, such as Halemaumau pit crater on the summit of Kilauea volcano and the paired lava lakes of the 1972 Mauna Ulu eruption (Decker, 1987; Tilling et al., 1987). Fire fountain eruptions in the lava lakes fed spatter to the rims, which were occasionally mantled by lava flows when the lava lakes overflowed their ramparts. The final eruptive phase was confined to the eastern vent, where fire fountains built ramparts on three sides of the vent that were not covered by subsequent lava flows, although lava from the eastern vent may have flowed into the western vent during lava highstands at this time (Shervais et al., 2002). Directions to Stop 2.4 Return to the frontage road and turn right; proceed SE along the frontage road (toward Mountain Home) for ≈5 mi (8 km) and turn right across the railroad tracks (you will be almost due north of Union Butte cinder cone). Continue south ~1.1 mi (1.8 km) to the Union Butte cinder cone.

B

C

Figure 10. Volcanic features near Mountain Home. (A) Plagioclase flotation cumulates in covered lava channel, Lockman Butte—note diktytaxitic texture near top of cumulate layer where exsolved gas has forced residual ferrobasaltic liquid from the interstices. (B) The Crater Rings, two large pit craters that fed lava lake eruptions and Peleanstyle fire fountains. (C) Pillow and hyaloclastite delta in basalt of Little Joe Butte (Strike Dam Road), which sits on lacustrine sediments of Glenns Ferry Formation (Lake Idaho); note water escape conduits, steam cavities, and southeast dip of foreset beds

Stop 2.4—Union Buttes: Holocene Cinder Cones, Late Alkaline Basalts Equivalent to Boise River Group 2 of Vetter and Shervais (1992) [N 43°9.565′, W 115°46.06′] Union Buttes are the two most prominent volcanic vents west of Mountain Home. Their stratigraphic position (overlying all older vents) and relative preservation suggest an age of <500,000 yr. The basalt of Union Buttes erupted from these two vents but was confined between the large Rattlesnake Springs shield volcano to the south and the smaller Crater Rings and Lockman Buttes vents to the west and north. The western Union Butte is larger than the eastern vent, and fed a small flow that flowed west toward Crater Rings. Both vents consist of cinder cones built on top of small shield volcanoes of similar composition basalt. The basalt of Union Buttes is distinct from almost all other basalts in the Mountain Home area because it contains large, clear phenocrysts of olivine but lacks plagioclase phenocrysts. The basalt of Union Buttes, and the similar basalt of Little Joe Butte (also known as the basalt of Strike Dam Road), are characterized by high K2O compared to other Mountain Home basalts,

Basaltic volcanism of the central and western Snake River Plain along with lower TiO2 and high-field strength trace element abundances. In terms of their chemical and age relations, they are similar to the Boise River Group 2 basalts of Vetter and Shervais (1992) and to the M3 basalts of White et al. (2002). The transition from high field strength-rich, alkali-poor tholeiites to potassium-rich, high field strength-poor transitional alkaline basalts around 600,000–700,000 yr ago is one of the most fundamental time-dependent transitions observed in the western Snake River Plain province, with the younger lavas characterized by OIB-like isotopic compositions. See Vetter and Shervais (1992) and White et al. (2002) for discussions. Directions to Stop 2.5 Continue south on gravel and paved roads ~2.6 mi (4.2 km) to intersection with Idaho Hwy 67. Turn right onto Hwy 67 and proceed SW toward Mountain Home Air Force Base. Bear right at the turn-off into the air base (away from the base) and continue SW on Hwy 67 for 6.5 mi (10.8 km) to a poorly marked intersection with Strike Dam Road. Turn left onto Strike Dam Road and proceed south for 6.75 mi (10.9 km) to the rim of the plateau; park at the top and walk down hill to the next stop. Stop 2.5—Basalt of Little Joe Butte: Subaerial Basalt on a Pillow Lava + Hyaloclastite Delta on Lake Idaho Sediments [N 42°57.648′, W 115°58.526′] The basalt of Little Joe Butte is olivine-phyric basalt that crops out along the western edge of the Mountain Home area (Cinder Cone Butte, Crater Rings SW quadrangles) and continues into adjacent quadrangles to the west (Little Joe Butte, Dorsey Butte) and south (Grand View, C.J. Strike Dam; Jenks et al., 1993). The source of this flow is Little Joe Butte, but it has also been mapped as the basalt of Strike Dam Road (Jenks et al., 1993). It consists of at least two major flow units, with a mappable flow front preserved in the upper unit, and appears to flow south along a former channel of the Canyon Creek drainage. Collapsed areas of this flow are commonly overlain surficially by “intermittent lake deposits,” i.e., dry lake beds (Shervais et al., 2002). This unit is largely subaerial, but its base is commonly pillowed along its southern margin, demonstrating the flow of subaerial lava into standing water of Lake Idaho. The subaqueous portion of the flow consists of pillow lava and hyaloclastite breccia forming foreset delta sequences with individual “rolled” pillows interspersed (Fig. 10C). The delta foresets dip to the SE, indicating the direction the lava flowed when it entered the lake. The subaqueous lava delta is ≈10 m thick in the bluffs overlooking C.J. Strike reservoir, where it flowed over flat-lying lake deposits—indicating the exact water depth in the lake at the time of eruption. Like the Bacon Butte Basalt at Stop 1.1B, the subaerial lava that overlies the pillow delta contains water escape conduits and giant vesicles (steam bubbles). The conduits and vesicles can be seen from the road but a close-up examination requires climbing up the pillow delta, which is inherently unstable—be careful if you go!

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Return to Hagerman Continue south on Strike Dam Road, across Snake River at Strike Dam Bridge, to Idaho Hwy 78. Drive E on Hwy 78 through Bruneau to Hammett. As time permits, we will make stops along this route to view the stratigraphy of the former Lake Idaho and the effects of the Bonneville flood. At Hammett, rejoin I-84 and return to motel. Day 3 (Partial Day) Twin Falls2 Directions to Stop 3.1 From Hagerman, drive north on Hwy 30 to Bliss, turn E (right) onto Hwy 26 to Shoshone. At main intersection in Shoshone, turn N (left) on Hwy 75. Drive N ~2.3 mi (3.7 km) and park on right. Hike E into the Shoshone lava flow to the lava channel. All of the stops for Day 3 are shown on Figure 11. Source vents for the basalt flows are indicated on Figure 12, a hillshade topographic map of the Twin Falls 30′ × 60′ quadrangle. Stop 3.1—Lava Channel is Shoshone Lava Flow; a‘a Basalt Fills Channel in Pahoehoe Flow [N 42°58.936′, W 114°18.22′] The Shoshone basalt flow (basalt of Black Butte Crater on the Idaho Geological Survey Twin Falls 30′ × 60′ geologic map) erupted from Black Butte Crater located in the Black Butte Crater quadrangle. The basalt is black, fine grained, vesicular to massive, and aphyric to olivine phyric. Olivine phenocrysts (Fo58–73) are visible in hand sample and are up to 2.0 mm in size. Most plagioclase (An31–72) displays a normal chemical zoning. Black Butte Crater is a broad shield volcano that rises 135 m above the surrounding topography. The volcano contains two large craters that are ~300 m across and 100 m in depth. This young lava field extends south to the Dietrich Butte quadrangle and then west into the Shoshone, Tunupa, and Gooding quadrangles, and terminates just west of the town of Gooding. Eruption of lava disrupted the confluence of the Big and Little Wood rivers, which was probably near Shoshone, and separated the two rivers to their present confluence at Gooding. Radiocarbon dating of underlying sediment baked by the flow (Kuntz et al., 1986) has yielded a date of 10,130 ± 350 yr B.P. for this basalt. Aside from a few areas covered by thin loess deposits or alluvium, most of the basalt is completely exposed, and almost all of the original flow morphology can be seen. The flow is dominated by a series of large lava tube and lava channel systems that can be seen on aerial photographs and topographic maps. Most of the Shoshone flow consists of smooth pavement outcrops of pahoehoe, overlain by sporadic a‘a flows; a‘a flows also fill lava channels and collapsed lava tubes. We plan to visit a large lava channel in the pahoehoe flow that is now filled with a‘a lava. This lava channel has levees that rise 60–70 ft above the surrounding lava, and the channel is as much as 70 ft deep (Fig. 13A).

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Stop 3.1 Stop 3.2 Stop 3.3 Day One Stops

Stop 3.4 Stop 3.5 Stop 3.6

Figure 11. Topographic map of central Snake River Plain showing Day 3 stops.

Figure 12 (opposite page). Hillshade of the Twin Falls area showing the names and locations of volcanic source buttes and the topographic expression of different basalt flows revealed through vertical exaggeration. The solid lines are main highways. The dashed line is the perimeter of the Twin Falls 30′ × 60′ quadrangle.

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A

inter-layered spatter and massive flows. Most outcrops along the flanks of the volcano can be found along terraced ridges concentric to the main crater. These terraces show the same interlayering of spatter and massive basalt, indicating that they may mark the position of older crater rims. The basalt of Crater Butte is dark gray, vesicular to massive, and fine grained, containing phenocrysts of plagioclase and olivine. Some of the samples taken from the massive layers display a slightly diktytaxitic texture but most are intergranular in texture. Olivine phenocrysts (Fo45–74) are small (<0.6 mm) and rounded. Plagioclase phenocrysts (An55–71) are generally 2.5 mm in size. Several large plagioclase phenocrysts display an oscillatory zoning, which may be due to pressure fluctuations during eruption or to magma mixing. Electron microprobe line scans across plagioclase phenocrysts reveal a strong normal chemical zonation. Directions to Stop 3.3 Return to paved road, turn right (S) and continue 3.3 mi (5.3 km) to Hwy 24. Turn right (W) onto Hwy 24 and continue to intersection at south end of Shoshone (≈8.3 mi/13.4 km). Turn left (S) onto Hwy 93 and drive 2.85 mi (4.6 km) to a gravel road on the left (east side) of road that leads to Notch Butte.

B Figure 13. Volcanic features of the central Snake River Plain north of Twin Falls, Idaho. (A) Lava channel in the Shoshone basalt. (B) Pit crater in Crater Butte.

Directions to Stop 3.2 Return to Shoshone and proceed NE on Hwy 93 toward Richfield. Turn right (S) ≈8.8 mi (14.2 km) from intersection in Shoshone onto a paved road. Drive south 2.5 mi (4.0 km) to the second dirt road on the right side of road. Turn right onto this dirt road and drive to the top of Crater Butte (~0.9 mi/1.5 km). Note: if we cannot get the vans up Crater Butte, we will drive to the top of nearby Dietrich Butte instead. Stop 3.2—Crater Butte [N 42°57.582′, W 114°16.724′] Crater Butte, located in the northeastern portion of the Dietrich quadrangle, is a steeply sloping shield volcano that rises ~140 m above the surrounding topography. The most striking feature of this volcano is the large 80 m deep crater in its center. This bowling pin-shaped crater is ~1300 m × 1000 m in size and trends in a NW–SE direction. The inner walls of the crater show thin layers of spatter interbedded with thicker beds of more massive basalt (Fig. 13B). The tops of the more massive beds are highly vesicular which indicates gaseous escape from lava channels beneath the cooler, more viscous upper crust. When the volcano was active, the crater probably contained a lava lake that would periodically spill over the spatter rim to create the

Stop 3.3—Notch Butte [N 42°53.036′, W 114°24.98′] Notch Butte is located in the southeastern corner of the Shoshone quadrangle and rises 110 m above the surrounding topography. The lava flows from this broad shield volcano cover 150 km2, including the southern half of the Shoshone quadrangle, west into the Tunupa and Gooding SE quadrangles, south into the Shoshone SW and SE quadrangles, and east into the Dietrich quadrangle. Several lobes flowed west to Gooding, continued south and west around Gooding Butte, and flowed into the ancestral Snake River canyon near Hagerman. The unit is equivalent to the Wendell Grade basalt of Malde et al. (1963), except for the canyon filling part at Hagerman, which they mapped as Sand Springs basalt. The unit has been renamed during recent mapping to follow the convention of naming flows after their source vent. The flows display a relatively young volcanic morphology with many flow features such as pressure ridges, collapsed lava tubes, and flow fronts still visible beneath a thin soil and loess layer. The loess mantle is mostly confined to depressions in the flow surfaces. Note the contrast between flows from this vent and those from the older Crater Butte vent and the younger Shoshone flow lavas. Mapping revealed three to four lobes of varying mineralogy within the flow field. In general, higher lobes closer to the vent were found to be rich in plagioclase, while those at lower elevations farther away from the vent contain less plagioclase and are rich in olivine. The higher lobes are interpreted to be younger, more fractionated, and therefore more viscous basalt (due to the plagioclase content), while those found at lower elevations are older, less fractionated, and less viscous. This difference in viscosity between the older and younger lava accounts for the fact

Basaltic volcanism of the central and western Snake River Plain that the older flows spread out over a large area away from the vent while the younger flows did not. The basalt of Notch Butte is a black, vesicular to massive, medium- to fine-grained, plagioclase- ± olivine-phyric basalt. Samples taken from some flow lobes contain large glomerocrysts of plagioclase and olivine, which are visible in hand sample. The basalt is intergranular to intersertal in texture and contains olivine phenocrysts (Fo35–73); large glomerocrysts of plagioclase (An59–71) and olivine are visible in thin section and are often up to 5 mm in size. The larger plagioclase phenocrysts are typically normally zoned. The view from the top of Notch Butte includes all of the major volcanoes in this part of the central Snake River plain: Crater Butte, Dietrich Butte, Owinza Butte, and Black Ridge Crater to the east; Lincoln Butte to the west; Bacon Butte, and Flat Top Butte to the south, and Wilson Butte and Rocky Butte to the southeast. From here, it is easy to contrast the relative ages of the flows, which can be estimated by the amount of cover indicated by farming. We can also see the western flow from Notch Butte, which extends some 40 km to the west where it flowed into the ancestral valley of the Snake River. At the time of Notch Butte eruptions, the ancestral Snake River was already at about its present position and elevation. Directions to Stop 3.4 Return to Hwy 93. Turn left (S) onto Hwy 93 and drive south 11.8 mi (19 km) to Hwy 25; turn left (E) onto Hwy 25 and after ≈0.6 mi (1 km) turn left (N) onto paved road; follow for 0.3 mi (0.5 km), then turn left (N) onto paved/gravel road that goes to top of Flat Top Butte (~1 mi/1.6 km from highway). Stop 3.4—Flat Top Butte [N 42°43.722′, W 114°24.75′] Flat Top Butte is an extremely large (≈450 km2) shield volcano located ~20 km north of Twin Falls in the Falls City 7.5′ topographic map quadrangle. The vent has a large (200 m), shallow depression at its summit that represents the former summit crater; the rim of this crater is now home to a forest of microwave transceivers. The flanks of the volcano are covered with a thick mantle of loess and soil that obscures the underlying basalt. This butte is farmed almost to its summit, showing that it is one of the older vents in this part of the plain. Radiometric ages for Flat Top Butte are in the 0.330–0.395 Ma range (Tauxe et al., 2004; Idaho Geological Survey, 2005, unpublished data). The basalt of Flat Top Butte flowed south and west filling the course of the ancestral Snake River as far as Thousand Springs, where Malde and Powers (1972) mapped the flows as the Thousand Springs Basalt. Flows from Flat Top Butte defined the present position of the Snake River by forcing the drainage southward against the regional north slope of the older shield volcanoes. Subsequently, the Snake River canyon was cut. From the summit area of Flat Top Butte we can look east toward the younger vents Wilson Butte, Rocky Butte, and Kimama Butte, north toward Bacon Butte and Notch Butte, and

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northeast toward Owinza Butte and Black Ridge Crater. Older buttes to the southeast include Hansen Butte, Skeleton Butte, Hazelton Butte, Milner Butte, and Burley Butte. Looking south toward the Snake River canyon and Twin Falls City, the track of Bonneville Flood overland flow was east to west obliquely across I-84. Along the canyon wall south of I-84 the overland floodwaters rejoined the component following the course of the river, forming cataracts in several places. Although the basalt of Rocky Butte typically has loess filling surface depressions, in the track of the floodwaters virtually all the loess is stripped away and patchy gravel deposits record local deposition during the flood. Directions to Stop 3.5 Return to Hwy 25; turn left (E) onto Hwy 25 and proceed 4.6 mi (7.4 km) to intersection with paved road. Turn S (right) and drive 5.4 mi (8.7 km) south—road goes from paved to gravel after ~2 mi (3.2 km). Continue on dirt road to SE around end of Wilson Butte flow lobe, ≈1 mi (1.6 km). Park anywhere. Stop 3.5—Wilson Butte Flow Lobe [N 42°37.82′, W 114°21.76′] Wilson Butte is a large shield volcano (≈10 km across). Its vent, composed of three small peaks, is rather unimpressive when compared to the estimated 250 km2 of lava that flowed down the slopes of the shield. The flows cover the majority of the Shoshone SE and half of the Star Lake quadrangles. The flow continues into the Shoshone SW and Hunt quadrangles, and it crosses Hwy 25 in the Twin Falls NE quadrangle and continues into the Kimberly quadrangle, possibly emptying into the Snake River around the Twin Falls rapids or Devil’s Corral area. The mode of the basalt is 35%–40% plagioclase, 10%–15% olivine, 10%–20% pyroxene, 5%–8% oxides, and 15%–20% glass. Plagioclase shows typical euhedral lath-shaped crystals with An content ranging from An66 to An36. Olivine ranges from Fo71 to Fo33. At this stop we will view a major flow lobe from Wilson Butte that stands 10–20 m above the surrounding lava fields and contains a giant complex of lava tubes that fed secondary flow lobes that flowed primarily west and south. The flow lobe is characterized by steep sides and a nearly flat upper surface that can be traced for tens of kilometers. This lava tube–channel system fed flows adjacent to the Snake River that have been mapped as Sand Springs basalt—in addition to flows from Rocky Butte (Hill 4526), which may dominate the Sand Springs unit. Wilson Butte and Rocky Butte have lavas that are nearly identical chemically, and both appear to have erupted at essentially the same time, although the Wilson Butte flow lobe appears to be younger than the adjacent flows from Rocky Butte. Tauxe et al. (2004) report radiometric age for “Sand Springs Basalt,” probably from Rocky Butte, as 0.095 Ma. The lava tube system is exposed in a series of windows that are open mainly on the eastern margin of the flow. Sitting on the Wilson Butte flow lobe here are coarse sediments deposited by the Bonneville flood. These sediments seem

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to represent a lag deposit from water flowing over or around the end of the lobe. At this elevation (~1173 m) the Bonneville Flood waters were relatively shallow. The relief of this flow lobe was just sufficient to divert most of the overland floodwater around the feature, but a few low spots allowed water to spill through and create small plunge pools and gravel deposits. The flood waters lost energy on the lee side of the flow lobe and deposited finer sediment that forms flat surfaces in broad depressions. Directions to Stop 3.6 Return to Hwy 25, turn W (left) and return to Hwy 93. Turn left onto Hwy 93 and drive south 8.6 mi (13.8 km) to Perrine Bridge in Twin Falls (you will cross I-84). Cross the bridge and enter turnout into viewpoint-tourist information area immediately on south side of bridge. Park in visitor center parking lot next to bridge. Stop 3.6—Perrine Bridge Overlook [N 42°35.8966′, W 114°27.266′] The Perrine Bridge viewpoint is one of the classic views of volcanic stratigraphy in the central Snake River Plain. The Snake River gorge is over 120 m deep here and contains two golf courses within the gorge just downstream from the bridge. Evil Kneivel’s famous attempt to jump the gorge on a rocket-powered motorcycle took place just upstream from here.

Sand Springs Basalt (Rocky or Wilson Buttes) Basalt of Flat Top Butte Hub Butte Basalt (?)

Glenns Ferry Fm.

Twin Falls Rhyolite

Figure 14. Volcanic stratigraphy at the Perrine Bridge, Twin Falls, Idaho. Dark rock at bottom of canyon is the 6.25 Ma Twin Falls rhyolites. The thickest section of basalt erupted from vents south of the river (Hub Butte, Sonnickson Butte), and is overlain by a thinner section of basalt from north of the current river (mostly Flat Top Butte with basalt of Rocky Butte and/or Wilson Butte) on top. Basalt and rhyolite are separated by thin wedges of clastic sediment in middle of section.

Looking north across the river we can see over 4 m.y. of volcanic history (Fig. 14). At the base of the section is the Twin Falls rhyolite (ca. 6.25 Ma; Armstrong et al., 1975) that forms the basement here, and which underlies Shoshone Falls east of Twin Falls. Sitting on the rhyolite (and on some patchy exposures of rhyolite breccia) are sediments that may correlate with the Pliocene Glenns Ferry Formation (Covington et al., 1985; Covington and Weaver, 1989). If this correlation holds, these sediments would represent lake and flood-plain sediments that were deposited in an extension of Lake Idaho. Basalts intercalated with these sediments just upstream from here have been dated at ca. 4 Ma by Armstrong et al. (1975). Sitting on the sediment and rhyolite is a thick section of basalt erupted from large vents to the south (Hub Butte and other southern buttes). These flows are overlain by flows from Hansen Butte and other volcanoes to the southeast, which pushed the river farther north and built the north slope of a large shield complex. Evidence downstream shows subsequent flows from Flat Top Butte burying the preexisting northward slope, presumably filling any previous east-west drainage, and reestablishing the course of the ancestral Snake River at the south edge of the Flat Top shield. By the time of eruptions from Wilson Butte and Rocky Butte, the Snake River had incised the older basalts along the margin of the Flat Top Butte shield. Evidence in the canyon wall shows local filling of the canyon by the younger flows. Subsequent flows from Flat Top Butte, Wilson Butte, and Rocky Butte followed the northern margin of these flows. The course of the Snake River then became established along the contact of the younger and older flows. The basalt of Flat Top Butte was formerly mapped as Thousand Springs Basalt by Malde and Powers (1972), while the overlying Sand Springs Basalt (Malde and Powers, 1972) is now known to represent flows from at least two vents in this area, Wilson Butte and Rocky Butte. However, valley-filling Sand Springs Basalt mapped along Cedar Draw south of the Snake River (Malde and Powers, 1972) is probably from Flat Top Butte. At this vantage point near Perrine Bridge, the Bonneville Flood at maximum discharge would have been mostly confined to the canyon except that overland floodwaters from the northeast were rejoining the canyon along here and the water depth exceeded the top of the canyon by ~15 m. Areas of cataracts and plunge pools on the north side of the canyon, such as the Blue Lake alcove, show the erosive power of the overland flow (Fig. 15). In the canyon bottom, the golf courses have been built on flood-stripped and molded surfaces of older basalt and rhyolite that have thin deposits of flood gravel and sand. The erosional and depositional features of the Bonneville Flood are time transgressive, albeit a short time period of only a few months. Unlike Glacial Lake Missoula that is thought to have emptied and produced a catastrophic flood duration of only several days, the catastrophic floodwater from Pleistocene Lake Bonneville was more complex (O’Connor, 1993). The early phases were in low discharge as the lake gently overtopped the divide at Red Rock Pass and it began to erode. As erosion accel-

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Figure 15. Topographic map of the area near Perrine Bridge showing features formed by the Bonneville Flood, including cataracts, plunge pools, scoured canyon walls, and giant gravel bars. Large arrows show inferred floodwater flow directions.

erated, the discharge grew to catastrophic proportions and was sustained for many weeks. However, as the level of Lake Bonneville dropped, gradually the discharge lowered. Ultimately, an equilibrium was achieved between the lake and the outlet, and the lake continued to drain into the Snake River drainage for at least hundreds of years. Given this backdrop, the overland floodwaters entering the canyon from the northeast represent flow during the greatest discharges, and would have ceased when the canyon could accommodate the flood. Features in the canyon, therefore, range from high-discharge rock scouring and deposition of giant, bouldery gravel bars, to lower-discharge thin sand and gravel deposits in the bottom of the canyon. Return to I-84. Enter freeway in east-bound direction (right turn). Return to Salt Lake City (approximate driving time: 4 hours). REFERENCES CITED Amini, M.H., Mehnert, H.H., and Obradovich, J.D., 1984, K-Ar ages of late Cenozoic basalts from the western Snake River Plain, Idaho: Isochron/ West, v. 41, p. 7–11. Anders, M.H., and Sleep, N.H., 1992, Magmatism and extension: the thermal and mechanical effects of the Yellowstone plume: Journal of Geophysical Research, v. 97, p. 15,379–15,393.

Armstrong, R.L., Leeman, W.P., and Malde, H.E., 1975, K-Ar dating, Quaternary and Neogene volcanic rocks of the Snake River Plain, Idaho: American Journal of Science, v. 275, p. 225–251. Armstrong, R.L., Harakal, J.E., and Neill, W.M., 1980, K-Ar dating of Snake River plain (Idaho) volcanic rocks; new results: Isochron/West, v. 27, p. 5–10. Bonnichsen, B., 1982a, The Bruneau-Jarbidge eruptive center, southwestern Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 237–254. Bonnichsen, B., 1982b, Rhyolite flows in the Bruneau-Jarbidge eruptive center, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 283–320. Bonnichsen, B., and Godchaux, M.M., 2002, Late Miocene, Pliocene, and Pleistocene geology of southwestern Idaho with emphasis on basalts in the Bruneau-Jarbidge, Twin Falls, and western Snake River Plain regions, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 233–312. Bonnichsen, B., and Kauffman, D.F., 1987, Physical features of rhyolite lava flows in the Snake River plain volcanic province, southwestern Idaho, in Fink, J.H., ed., The Emplacement of Silicic Domes and Lava Flows: Geological Society of America Special Paper 212, p. 119–145. Bonnichsen, B., White, C.M., and McCurry, M., eds., 2002, Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, 482 p. Braile, L.W., Smith, R.B., Ansorge, J., Baker, M.R., Sparlin, M.A., Prodehl, C., Schilly, M.M., Healy, J.H., Meuller, S., and Olsen, K.H., 1982, The Yellowstone–Snake River Plain seismic profiling experiment: Crustal structure of the eastern Snake River Plain: Journal of Geophysical Research, v. 87, p. 2597–2609.

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Camp, V.E., 1995, Mid-Miocene propagation of the Yellowstone mantle plume head beneath the Columbia River Basalt source region: Geology, v. 23, p. 435–438, doi: 10.1130/0091-7613(1995)023<0435:MMPOTY>2.3.CO;2. Camp, V.E., and Ross, M.E., 2004, Mantle dynamics and genesis of mafic magmatism in the intermontane Pacific Northwest: Journal of Geophysical Research, v. 109, B08204, doi: 10.1029/2003JB002838. Cecil, L.D., Welhan, J.A., Green, J.R., Frape, S.K., and Sudicky, E.R., 2000, Use of chlorine-36 to determine regional-scale aquifer dispersivity, eastern Snake River Plain aquifer, Idaho: Nuclear instruments & methods in physics research, section B, Beam interactions with materials and atoms, v. 172, p. 679–687, doi: 10.1016/S0168-583X(00)00216-0. Christiansen, R.L., 1982, Late Cenozoic volcanism of the Island Park area, eastern Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 345–368. Christiansen, R.L., Foulger, G.R., and Evans, J.R., 2002, Upper-mantle origin of the Yellowstone hotspot: Geological Society of America Bulletin, v. 114, no. 10, p. 1245–1256, doi: 10.1130/0016-7606(2002)114<1245: UMOOTY>2.0.CO;2. Clemens, D.M., and Wood, S.H., 1993, Late Cenozoic volcanic stratigraphy and geochronology of the Mount Bennett Hills, central Snake River plain, Idaho: Isochron/West, v. 60, p. 3–14. Cooke, M.F., 1999, Geochemistry, Volcanic stratigraphy, and hydrology of Neogene basalts, central Snake River Plain, Idaho [M.S. thesis]: Columbia, South Carolina, University of South Carolina, 125 p. Cooke, M.F., and Shervais, J.W., 1999, Stratigraphic controls of basaltic volcanism on groundwater recharge and conductivity in the central Snake River Plain, Idaho: Geological Society of America Abstracts with Programs, v. 31, no. 4, p. A8. Craig, H., 1997, Helium isotope ratios in Yellowstone Park and along the Snake River plain; backtracking the Yellowstone Hotspot: Eos (Transactions American Geophysical Union), v. 78, p. 801. Covington, H.R., and Weaver, J.N., 1989, Geologic map of the profile of the northern wall of Snake River canyon: U.S. Geological Survey Map I1947, A–E, scale 1:24000. Covington, H.R., Whitehead, R.L., and Weaver, J.N., 1985, Ancestral canyons of the Snake River: Geology and hydrology of canyon-fill deposits in the Thousand Springs area, south-central Snake River Plain, Idaho: Boise, Idaho, Geological Society of America, Rocky Mountain Section, April 1985, Guide Book, 30 p. Decker, R.W., 1987, Dynamics of Hawaiian volcanoes: An overview; in Decker, R.W., Wright, T.L., and Stauffer, P.H., eds., Volcanism in Hawaii: U.S. Geological Survey Professional Paper 1350, v. 2, p. 997–1018. Doe, B.R., Leeman, W.P., Christiansen, R.L., and Hedge, C.E., 1982, Lead and strontium isotopes and related trace elements as genetic tracers in the upper Cenozoic rhyolite-basalt association of the Yellowstone Plateau volcanic field: Journal of Geophysical Research, v. 87, p. 4785–4806. Draper, D.S., 1991, Late Cenozoic bimodal magmatism in the northern Basin and Range province of southeastern Oregon: Journal of Volcanology and Geothermal Research, v. 47, p. 299–328, doi: 10.1016/03770273(91)90006-L. Dueker, K., and Humphreys, E., 1990, Upper mantle velocity structure of the Great Basin: Geophysical Research Letters, v. 17, no. 9, p. 1327–1330. Dueker, K.G., Schutt, D.L., Yuan, H., and Fee, D., 2004, New seismic constraints for the Yellowstone hotspot: Eos (Transactions American Geophysical Union), v. 85/47, Fall Meeting Supplement, Abstract 51B-0554. Farnetani, C.G., and Samuel, H., 2004, Dynamics of thermochemical plumes: Eos (Transactions American Geophysical Union), v. 85/47, Fall Meeting Supplement, Abstract 44B-03. Geist, D.J., and Richards, M., 1993, Origin of the Columbia River plateau and Snake River Plain: deflection of the Yellowstone plume: Geology, v. 21, p. 789–792, doi: 10.1130/0091-7613(1993)021<0789:OOTCPA>2.3.CO;2. Gillerman, V.S., 2004, Diversity in the Banbury Basalt: hydrovolcanoes, sediments and structures of the Banbury and Thousand Springs area, Snake River canyon, Idaho: Geological Society of America Abstracts with Programs, v. 34, no. 6, p. 86. Godchaux, M.M., and Bonnichsen, B., 2002, syneruptive magma-water and posteruptive lava-water interactions in the western Snake River Plain, Idaho, during the past 12 million years, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 387–434.

Godchaux, M.M., Bonnichsen, B., and Jenks, M.D., 1992, Types of phreatomagmatic volcanoes in the western Snake River Plain, Idaho, USA: Journal of Volcanology and Geothermal Research, v. 52, p. 1–25, doi: 10.1016/0377-0273(92)90130-6. Greeley, R., 1982, The style of basaltic volcanism in the eastern Snake River Plain, Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 407–422. Hart, W.K., and Aronson, J.L., 1983, K-Ar ages of rhyolites from the western Snake River Plain area, Oregon, Idaho, and Nevada: Isochron/West, v. 36, p. 17–19. Honjo, N., McElwee, K.R., Duncan, R.A., and Leeman, W.P., 1986, K-Ar ages of volcanic rocks from the Magic Reservoir eruptive center, Snake River plain, Idaho: Isochron/West, v. 46, p. 15–17. Honjo, N., Bonnichsen, B., Leeman, W.P., and Stormer, J.C., Jr., 1992, Mineralogy and geothermometry of high-temperature rhyolites from the central and western Snake River plain: Bulletin of Volcanology, v. 54, no. 3, p. 220–237. Howard, K.A., and Shervais, J.W., 1973, Geologic map of Smith Prairie, Elmore County, Idaho: U.S. Geological Survey Map I-818, scale 1:24,000. Howard, K.A., Shervais, J.W., and McKee, E.H., 1982, Canyon-filling lavas and lava dams on the Boise River, Idaho, and their significance for evaluating downcutting during the last two million years, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 629–641. Humphreys, E.D., and Dueker, K.G., 1994, Western U.S. upper mantle structure: Journal of Geophysical Research, B, Solid Earth and Planets, v. 99, no. 5, p. 9615–9634, doi: 10.1029/93JB01724. Humphreys, E.D., Dueker, K.G., Schutt, D.L., and Smith, R.B., 2000, Beneath Yellowstone; evaluating plume and nonplume models using teleseismic images of the upper mantle: GSA Today, v. 10, no. 12, p. 1–7. Hughes, S.S., Smith, R.P., Hackett, W.R., and Anderson, S.R., 1999, Mafic volcanism and environmental geology of the eastern Snake River plain, Idaho, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the Geology of Eastern Idaho: Idaho Museum of Natural History, p. 143–168. Hughes, S.S., McCurry, M., and Geist, D.J., 2002, Geochemical correlations and implications for the magmatic evolution of basalt flow groups at the Idaho National Engineering and Environmental Laboratory, in Link, P.K., and Mink, L.L., eds., Geology, hydrogeology, and environmental remediation; Idaho National Engineering and Environmental Laboratory, eastern Snake River plain, Idaho: Geological Society of America Special Paper 353, p. 151–173. Iyer, H.M., 1984, A review of crust and upper mantle structure studies of the Snake River Plain-Yellowstone volcanic system: a major lithospheric anomaly in the western USA: Tectonophysics, v. 105, p. 291–308, doi: 10.1016/0040-1951(84)90209-9. Jenks, M.D., and Bonnichsen, B., 1989, Subaqueous basalt eruptions into Pliocene Lake Idaho, Snake River plain, Idaho, in Chamberlin, V.E., Breckinridge, R.M., and Bonnichsen, B., eds., Guidebook of the Geology of Northern and Western Idaho and Surrounding Areas: Idaho Geological Survey Bulletin 28, p. 17–34. Jenks, M.D., Bonnichsen, B., and Godchaux, M.M., 1993, Geologic maps of the Grand View-Bruneau area, Owyhee County, Idaho: Idaho Geological Survey Technical Report 93-2, 21 p., scale 1:24,000. Jordan, M., Smith, R.B., and Waite, G.P., 2004, Tomographic Images of the Yellowstone Hotspot Structure: Eos (Transactions American Geophysical Union), v. 85/47, Fall Meeting Supplement, Abstract 51B-0556. Kauffman, D.F., and Bonnichsen, B., 1990, Geologic map of the Little Jacks Creek, Big Jacks Creek, and Duncan Creek wilderness study areas, Owyhee County, Idaho: U.S. Geological Survey Miscellaneous Field Studies Map MF-2142, scale 1: 50,000. Kimmel, P.G., 1982, Stratigraphy, age, and tectonic setting of the MiocenePliocene lacustrine sediments of the western Snake River plain, Oregon and Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 559–558. King, S.D., 2004, Where plumes live: Eos (Transactions, American Geophysical Union), v. 85/47, Fall Meeting Supplement, Abstract 44B-01. Kuntz, M.A., Champion, D.E., Lefebvre, R.H., and Covington, H.R., 1988, Geologic map of the Craters of the Moon, Kings Bowl, and Wapi lava fields and the Great Rift volcanic rift zone, south-central Idaho: U.S. Geological Survey Miscellaneous Investigations Series Map I-1632, scale 1:100,000.

Basaltic volcanism of the central and western Snake River Plain Kuntz, M.A., Champion, D.E., Spiker, E.C., Lefebvre, R.H., and McBroome, L.A., 1982, The Great Rift and the evolution of the Craters of the Moon Lava Field, Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 423–437. Kuntz, M.A., Spiker, E.C., Rubin, M., Champion, D.E., and Lefebvre, R.H., 1986, Radiocarbon studies of Holocene-latest Pleistocene lava flows of the Snake River Plain, Idaho: data, lessons, interpretations: Quaternary Research, v. 25, p. 163–176, doi: 10.1016/0033-5894(86)90054-2. Kuntz, M.A., Covington, H.R., and Schorr, L.J., 1992, An overview of basaltic volcanism of the eastern Snake River Plain, Idaho, in Link, P.K., Kuntz, M.A., and Platt, L.B., eds., Regional Geology of Eastern Idaho and Western Wyoming: Geological Society of America Memoir 179, p. 227–267. Leeman, W.P., 1982, Development of the Snake River Plain–Yellowstone Plateau province, Idaho and Wyoming: An overview and petrologic model, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 155–177. Lindholm, G.F., and Vaccaro, J.J., 1988, Region 2, Columbia Lava Plateau, in Back, W., Rosenshein, J.S., and Seabar, P.R., eds., Hydrogeology, Geology of North America: Geological Society of North America, v. O-2, p. 37–50. Link, P.K., and Fanning, C.M., 1999, Late Miocene Snake River flowed south into the Humboldt drainage: detrital zircon evidence: Geological Society of America Abstracts with Programs, v. 31, no. 4, p. A22. Mabey, D.R., 1976, Interpretation of a gravity profile across the western Snake River Plain, Idaho: Geology, v. 4, p. 53–55, doi: 10.1130/00917613(1976)4<53:IOAGPA>2.0.CO;2. Mabey, D.R., 1978, Regional gravity and magnetic anomalies in the eastern Snake River Plain, Idaho: U.S: Geological Survey Journal of Research, v. 6, no. 5, p. 553–562. Mabey, D.R., 1982, Geophysics and tectonics of the Snake River Plain, Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 139–153. Malde, H.E., 1968, The catastrophic late Pleistocene Bonneville Flood in the Snake River Plain, Idaho: U.S. Geological Survey Professional Paper 596, 52 p. Malde, H.E., 1991, Quaternary geologic and structural history of the Snake River Plain, Idaho and Oregon, in Morrison, R.B., ed., Quaternary Non-glacial Geology: Conterminous United States: Boulder, Colorado, Geological Society of America, The Decade of North American Geology, v. K-2, p. 251–281. Malde, H.E., and Powers, H.A., 1962, Upper Cenozoic stratigraphy of western Snake River Plain, Idaho: Geological Society of America Bulletin, v. 73, p. 1197–1220. Malde, H.E., and Powers, H.A., 1972, Geologic map of the Glenns Ferry-Hagerman area, west-central Snake River Plain, Idaho: U.S. Geological Survey Miscellaneous Investigations Map I-696, scale 1:48,000, 2 sheets. Malde, H.E., Powers, H.A., and Marshall, C.H., 1963, Reconnaissance geologic map of west-central Snake River Plain, Idaho: U.S. Geological Survey Miscellaneous Geological Investigations Map I-373, scale 1:125,000. Matthews, S.M., 2000, Geology of Owinza Butte, Shoshone SE, and Star Lake quadrangles: Snake River Plain, southern Idaho [M.S. thesis]: Columbia, South Carolina, University of South Carolina, 110 p. McCurry, M., and Hackett, W.R., 1999, Genesis of Quaternary rhyolites in Southeast Idaho; implications for the Yellowstone–Snake River plain hotspot system: Geological Society of America Abstracts with Programs, v. 31, no. 4, p. 24. McGee, J., and Shervais, J.W., 1997, Flotation cumulate in a Snake River Plain ferrobasalt: Petrologic study of a possible lunar analogue: Geological Society of America Abstracts with Programs, v. 29, no. 6, p. A136. McQuarrie, N., and Rodgers, D.W., 1998, Subsidence of a volcanic basin by flexure and lower crustal flow; the eastern Snake River plain, Idaho: Tectonics, v. 17, no. 2, p. 203–220, doi: 10.1029/97TC03762. Montelli, R., Nolet, G., Dahlen, F.A., Masters, G., Engdahl, R., and Hung, S.H., 2003, Finite-frequency tomography reveals a variety of plumes in the mantle: Science, v. 303, p. 338–343, doi: 10.1126/science.1092485. Morgan, L.A., 1992, Stratigraphic relations and paleomagnetic and geochemical correlations of ignimbrites of the Heise volcanic field, eastern Snake River Plain, eastern Idaho and western Wyoming, in Link, P.K., Kuntz, M.A., and Platt, L.B., eds., Regional Geology of Eastern Idaho and Western Wyoming: Geological Society of America Memoir 179, p. 215–226. Morgan, W.J., 1972, Plate motions and deep mantle convection, in Shagam, R., et al., eds., Studies in earth and space sciences: Geological Society of America Memoir 132, p. 7–22.

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O’Connor, J.E., 1993, Hydrology, hydraulics, and geomorphology of the Bonneville Flood: Geological Society of America Special Paper 274, 83 p. Peng, X., and Humphreys, E.D., 1998, Crustal velocity structure across the eastern Snake River plain and the Yellowstone Swell: Journal of Geophysical Research, B, Solid Earth and Planets, v. 103, no. 4, p. 7171–7186, doi: 10.1029/97JB03615. Pierce, K.L., and Morgan, L.A., 1992, The track of the Yellowstone hot spot: volcanism, faulting, and uplift, in Link, P.K., Kuntz, M.A., and Platt, L.B., eds., Regional Geology of Eastern Idaho and Western Wyoming: Geological Society of America Memoir 179, p. 1–53. Pierce, K.L., Morgan, L.A., and Saltus, R.W., 2002, Yellowstone plume head: postulated tectonic relations to the Vancouver slab, continental boundaries, and climate, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 5–33. Priestley, K.F., and Orcutt, J., 1982, Extremal travel time inversion of explosion seismology data from the eastern Snake River plain, Idaho, YellowstoneSnake River plain symposium: Journal of Geophysical Research, v. B, p. 2634–2642. Repenning, C.A., Weasma, T.R., and Scott, G.R., 1995, The early Pleistocene (latest Blancan-earliest Irvingtonian) Froman Ferry fauna and history of the Glenns Ferry Formation, southwestern Idaho: U.S. Geological Survey Bulletin 2105, 86 p. Rodgers, D.W., Ore, H.T., Bobo, R.T., McQuarrie, N., and Zentner, N., 2002, Extension and subsidence of the eastern Snake River Plain, Idaho, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 121–155. Saltzer, R.L., and Humphreys, E.D., 1997, Upper mantle P-wave velocity structure of the eastern Snake River Plain and its relationship to geodynamic models of the region: Journal of Geophysical Research, B, Solid Earth and Planets, v. 102, no. 6, p. 11829–11841, doi: 10.1029/97JB00211. Shervais, J.W., Shroff, G., Vetter, S.K., Matthews, S., Hanan, B.B., and McGee, J.J., 2002, Origin of the western Snake River Plain: Implications from stratigraphy, faulting, and the geochemistry of basalts near Mountain Home, Idaho, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and Magmatic Evolution of the Snake River Plain Volcanic Province: Idaho Geological Survey Bulletin 30, p. 343–361. Shervais, J.W., Vetter, S.K., and Hanan, B.B., 2004, Basaltic Volcanism of the Central Snake River Plain, Idaho: Geological Society of America Abstracts with Programs, v. 36, no. 4, p. 98. Smith, G.R., and Stearley, R.F., 1999, Fish paleoecology and late Cenozoic history of the Snake River Plain: Geological Society of America Abstracts with Programs, v. 31, no. 4, p. A56. Smith, G.R., Swirydczuk, K., Kimmel, P.G., and Wilkinson, B.H., 1982, Fish biostratigraphy of late Miocene to Pleistocene sediments of the western Snake River Plain, Idaho, in Bonnichsen, B., and Breckenridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 519–542. Smith, R.B., and Braile, L.W., 1994, The Yellowstone hotspot: Journal of Volcanology and Geothermal Research, v. 61, p. 121–187, doi: 10.1016/03770273(94)90002-7. Stearns, H.T., Crandall, L., and Steward, W.G., 1938, Geology and groundwater resources of the Snake River Plain in southeastern Idaho: U.S. Geological Survey Water-Supply Paper 774, 268 p. Suppe, J., Powell, C., and Berry, R., 1975, Regional topography, seismicity, Quaternary volcanism, and the present day tectonics of the western United States: American Journal of Science, v. 275A, p. 397–436. Tauxe, L., Luskin, C., Selkin, P., Gans, P., and Calvert, A., 2004, Paleomagnetic results from the Snake River Plain: contribution to the time-averaged field global database: Geochemistry Geophysics Geosystems (G3), v. 5, no. 8, QH13. Tilling, R.I., Wright, T.L., and Millard, H.T., Jr., 1987, Trace element chemistry of Kilauea and Mauna Loa lava in space and time: a reconnaissance, in Decker, R.W., Wright, T.L., and Stauffer, P.H., eds., Volcanism in Hawaii: U.S. Geological Survey Professional Paper 1350, v. 1, p. 641–690. Vetter, S.K., and Shervais, J.W., 1992, Continental basalts of the Boise River Group near Smith Prairie, Idaho: Journal of Geophysical Research, B, Solid Earth and Planets, v. 97, no. 6, p. 9043–9061. Vetter, S.K., and Shervais, J.W., 1997, Basaltic volcanism of the Bruneau-Jarbidge eruptive center, southwest, Idaho: Geological Society of America Abstracts with Programs, v. 29, no. 6, p. A298.

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Printed in the USA

Geological Society of America Field Guide 6 2005

From cirques to canyon cutting: New Quaternary research in the Uinta Mountains Jeffrey S. Munroe Geology Department, Middlebury College, Middlebury, Vermont 05753, USA Benjamin J.C. Laabs Geology Department, Gustavus Adolphus College, St. Peter, Minnesota 56082, USA Joel L. Pederson Geology Department, Utah State University, Logan, Utah 84322, USA Eric C. Carson Geology Department, San Jacinto College, Houston, Texas 77049, USA

ABSTRACT The Quaternary record of the Uinta Mountains of northeastern Utah has been studied extensively over the past decade, improving our understanding of the Pleistocene glacial record and fluvial system evolution in a previously understudied part of the Rocky Mountains. Glacial geomorphology throughout the Uintas has been mapped in detail and interpreted with reference to other well-studied localities in the region. In addition, studies in Browns Park and Little Hole in the northeastern part of the range have provided information about paleoflooding, canyon cutting, and integration of the Green River over the Uinta Mountain uplift. Notable contributions of these studies include (1) constraints on the timing of the local last glacial maximum in the southwestern Uintas based on cosmogenic surface exposure dating, (2) insight into the relationship between ice dynamics and bedrock structure on the northern side of the range, and (3) quantification of Quaternary incision rates along the Green River. This guide describes a circumnavigation of the Uintas, visiting particularly well-documented sites on the north and south flanks of the range and along the Green River at the eastern end. Keywords: Uinta Mountains, Colorado Plateau, Last Glacial Maximum, Quaternary. INTRODUCTION

on aspects of these records in the range. Their work was unified by recognition of the characteristics that make the Uinta Mountains important to our overall understanding of landscape evolution in the interior western United States. First, the Uintas contain an unusually complete record of post-Laramide tectonics, erosion, and drainage integration, and they contained an extensive glacier complex during the Pleistocene glaciations. Second, the unique

For much of the past century, the Quaternary record of the Uinta Mountains escaped the attention focused on the neighboring Colorado Front Range, Wind River Range, and Yellowstone Plateau. That situation changed in the mid-1990s when researchers interested in glacial and fluvial geomorphology began working

Munroe, J.S., Laabs, B.J.C., Pederson, J.L., and Carson, E.C., 2005, From cirques to canyon cutting: New Quaternary research in the Uinta Mountains, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 53–78, doi: 10.1130/2005.fld006(03). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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east-west orientation of the mountains allows study of paleoprecipitation gradients in a direction parallel to primary storm tracks and moisture transport. Third, the location and orientation of the range allow the Quaternary record in the Uintas to function as a link between the middle Rockies and Colorado Plateau to the Great Basin and the Sierra Nevada. Finally, numerous famous geologists worked in the Uintas in the late 1800s and early 1900s, including Hayden (1871), Powell (1876), King (1878), Atwood (1909), and Bradley (1936). Many of their seminal ideas regarding tectonics and landscape evolution evolved through consideration of the evidence observed in the Uinta Mountains. Trip Overview This field trip provides an overview of recent work illuminating aspects of the glacial and fluvial history of the Uinta Mountains from the Last Glacial Maximum (MIS-2, ca. 22–18 ka) to

the present. Much of this work was conducted within the 456,704acre High Uintas Wilderness Area, which is jointly administered by the Ashley and Wasatch-Cache National Forests. Depending upon the season, it is often not possible to visit sites in the interior of the Uintas. However, this trip takes full advantage of viewpoints and exposures around the perimeter of the range in order to summarize our work and to introduce lingering questions. The route of this trip circumnavigates the Uinta Mountains in a counter-clockwise direction (Fig. 1). The first day focuses on glacial deposits on the south slope of the range, where county roads (CR) and Forest Service roads (FR) allow access to glacial deposits in the Lake Fork and Yellowstone River valleys. The second day of the trip explores the post-Laramide tectonic and drainage evolution of the eastern end of the range, including the Quaternary stratigraphy of the Green River and recognition of a significant paleoflood event. The final day continues westward along the north slope of the Uintas, visiting late Quaternary gla-

Figure 1. Route of the field trip through northeastern Utah and southwestern Wyoming. Inset shows the state of Utah, the extent of Lake Bonneville at the Last Glacial Maximum (LGM), the Uinta Mountains (stippled pattern on inset), and the field trip route. Larger map shows the reconstructed outlines of LGM glaciers (dark gray), the route of the trip (solid white line), alternate roads in case of early season snow (dotted white lines), and numbered stops (corresponding to text). BP—Browns Park; RCL—Red Canyon Lodge.

New Quaternary research in the Uinta Mountains cial and fluvial deposits in several localities before returning to Salt Lake City. As a note to guidebook users, mechanized equipment of any kind is prohibited within federally designated wilderness areas. Thus, access to field sites within the High Uintas Wilderness is by foot or by horseback only. Limitations also exist on the number of people and heads of stock allowed per party. Please check with either the Ashley National Forest (+1-435-789-1181, Vernal, Utah) or the Wasatch-Cache National Forest (+1-307-789-3194, Evanston, Wyoming) for details if you are interested in visiting the Uinta backcountry. Physical Setting The Uinta Mountains are the longest east-west–trending mountain range in the conterminous United States, extending ~200 km eastward from the Wasatch Mountains at Kamas, Utah, into northwestern Colorado. Physiographically, the range can be divided into two sections: the western glaciated Uintas and the eastern nonglaciated Uintas (Hansen, 1986). The boundary between the two subprovinces is located ~40 km west of the Utah-Colorado border, along a line extending north from Vernal (Fig. 1). The core of the western Uintas, centered on Kings Peak, the highest mountain in Utah at 4136 m, is referred to as the “High Uintas.” The drainage systems in the Uinta Mountains flow primarily north and south away from the crest of the range. On the north flank, short tributaries originating in alpine cirques drain into main streams in the glaciated valleys. Bear River, at the western end of the range, drains into Great Salt Lake; all other streams flow into the Green River upstream from Flaming Gorge dam. On the south flank, tributaries in large compound cirques feed into deep, narrow canyons that descend to the Uinta Basin. All streams on the south flank either flow into the Duchesne River (a tributary of the Green River) or directly into the Green River. Bedrock Geology The Uinta Mountains are cored by a series of Precambrian orthoquartzite, shale, and sandstone units known as the Uinta Mountain Group (UMG) (Bryant, 1992; Bradley, 1995; Hansen, 1969; Ritzma, 1959; Wallace and Crittenden, 1969; Dehler et al., 2006). These sediments, with a total thickness in excess of 7 km (Hansen, 1965; Stone, 1993), accumulated in an intracratonic rift basin near the suture of the Wyoming Archean craton (the Cheyenne Belt) and Paleoproterozoic terrane (the Colorado Province) ca. 800 Ma (Condie et al., 2001). The uppermost unit of the Uinta Mountain Group is the Red Pine Shale, which is unconformably overlain by thin (≤50-m-thick) sandstone of the Cambrian Tintic and Lodore Formations in parts of the western and eastern Uintas, but is more commonly overlain directly by Mississippian Madison Limestone. The Madison Limestone is locally >200 m thick (Bryant, 1992), forms hogbacks in some places where it is dissected by glacial valleys, and forms karst

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topography at ~3000 m in the headwaters of the Rock Creek and Duchesne River. Stratigraphically above this unit are the Mississippian Humbug (limestone and dolomite) and Doughnut (shale, limestone, and dolomite) Formations, which are 10–50 m thick. Large landslides are common near the contact of these two units, one of which can be observed at Stop 1.2. Upper Paleozoic strata include the Permian Weber Sandstone, which forms local hogbacks and which is overlain by upper Permian and Mesozoic strata of varying thickness at the mouths of glacial valleys. Flat-lying, weakly cemented Tertiary gravel deposited during uplift of the Uinta Mountains is generally found above angular unconformities with tilted, pre-Cenozoic bedrock. On the north slope, these deposits are mapped as the extensive Wasatch Formation (Bryant, 1992). On the south flank, these gravels include fluvial sandstone, conglomerate, and colluvial or mudflow diamictite of the Duchesne River Formation, which are most common on unglaciated uplands between the Duchesne River and Whiterocks drainage basins (Bryant, 1992). Sandstone and conglomerate beds in this unit are poorly indurated and are a source for mass-wasting and large alluvial fan deposits in tributaries of glacial valleys. The age of this unit is considered to be early Tertiary; however, recent work by D. Sprinkel (2004, personal commun.) suggests that several gravels currently mapped as part of this unit may be significantly younger (late Tertiary or early Pleistocene). Igneous rocks are almost entirely absent from the Uintas. Several small mafic dikes exist along the crest of the uplift (Ritzma, 1983). Dates from these dikes range from 552 ± 17 to 453 ± 29 Ma on the basis of Rb/Sr wholerock and K/Ar methods (Hansen et al., 1982; Ritzma, 1983; Rowley et al., 1985). Structural Geology Structurally, the Uintas are the surface expression of a broad doubly plunging anticlinal uplift, the axis of which is roughly concordant with the ridge crest. The axis of the uplift extends westward through the Wasatch Range and emerges as the Cottonwood Uplift at the Wasatch Front (Butler et al., 1920; Hansen, 1986; Bradley, 1995; Paulsen and Marshak, 1999). At the eastern end, the axis extends through northwestern Colorado and merges with the White River uplift (Tweto, 1976; Hansen, 1986). The Uinta anticline is asymmetrical, with its axis much closer to the north flank of the fold. Consistent with this geometry, the Precambrian and Paleozoic strata in the Uinta Mountains are subhorizontal at the crest of the range, gently dipping to the south on the southern flank and steeply dipping to the north on the northern flank. Uplift of the Uinta Mountains occurred during the Laramide Orogeny (late Mesozoic to early Cenozoic time) and was characterized by folding and thrusting along a system of east-west– trending faults on the north and south flanks of the range. Folding of the Uinta arch accommodated deformation by crustal shortening, and thrust faulting accommodated brittle deformation on the north and south flanks (Bradley, 1995). North-dipping thrust

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faults on the south flank of the Uintas are now buried by sediments in the Uinta Basin; however, south-dipping thrust faults are well-exposed on the north flank (Bradley, 1995). Displacement also occurred to a lesser extent on normal faults during the later stages of the Laramide Orogeny; one prominent normal fault is the South Flank fault, which juxtaposed Madison Limestone and quartzite of the upper UMG (Bryant, 1992). Once Laramide uplift ceased, there was an episode of tectonic stability throughout the Rocky Mountains during which significant pedimentation took place (Epis and Chapin, 1975; Mears, 1993). This includes the extensive Gilbert Peak erosion surface developed around the flanks of the Uintas (Sears, 1924; Bradley, 1936; Hansen 1986). Two Tertiary deposits record relatively subtle sedimentation and deformation in the Uintas: the Oligocene Bishop Conglomerate, which lies atop the Gilbert Peak erosion surface, and the Miocene Browns Park Formation. The regional tilting of the Bishop Conglomerate–Gilbert Peak surface and faulting associated with the Browns Park graben record the collapse of the eastern part of the Uinta uplift in late Oligocene–Miocene time (Bradley, 1936; Hansen, 1984; 1986). Geomorphology Although glaciers are absent from the modern Uinta Mountains, the spectacular alpine landscape of the High Uintas testifies to extensive Pleistocene glaciation. Cirque glaciers coalesced to form confined valley glaciers that apparently never surmounted the crest of the range or the major drainage divides, except in a few drainages at the west end of the range. As a result, the topography of the Uinta Mountains is dominated by deep, glacially scoured valleys separated by broad, unglaciated interfluves. Terminal moraines dating to the Last Glacial Maximum and older glaciations are present in most glaciated valleys, and outwash valley trains and meltwater channels are locally well developed. At the highest elevations, abundant and diverse periglacial deposits cover many slopes and the tops of most ridges and peaks of the Uintas. These include talus, protalus ramparts, felsenmeer, solifluction features, patterned ground, nivation hollows, and rock glaciers. The ages of these deposits are controversial. Grogger (1974) proposed that current periglacial activity is only capable of locally modifying older deposits. However, apparently active nonsorted circles are present at higher elevations. Furthermore, many rock glaciers have features suggestive of motion, including fresh unstable frontal slopes steeper than the angle-of-repose, transverse and longitudinal surface furrows, springs discharging 0 °C water in late summer, and surface meltwater ponds (Wahrhaftig and Cox, 1959). On the basis of these characteristics, Osborn (1973) reported that several rock glaciers in the western Uintas appear to be active. It is also worth noting that Bauer (1985) reported the presence of perennial ice in talus at an elevation of 2500 m in the western Uintas, strongly suggesting that active rock glaciers and permafrost are a possibility at higher elevations. At lower elevations, Holocene alluvial deposits mantle valley floors and merge with a series of outwash terraces and moraines

where streams flow into the Green River and Uinta basins. The stream planforms are dictated by the glacial history. In headwater regions, streams originate from talus slopes and cirque lakes and feed directly into mainstem channels. The interfluves and steep valley walls are drained by distributed surface flow and short tributaries, often with summer discharges of no more than ~0.05 m3/s. The mainstem channels are typically confined by the valley walls in areas where resistant Precambrian and Paleozoic bedrock outcrops. The stream channel gradients through these areas are steep (water surface slopes in excess of 0.002 m/m), with the bed comprised primarily of cobbles and boulders. In areas where less resistant bedrock is exposed, the valley floors widen appreciably into low-gradient meadows locally more than 300 m wide. Channel bed sediment in these meadows is dominated by sand and silt in point bars and sand and gravel on channel bottoms, and slopes are commonly <0.0005 m/m. History of Research Very little previous nonglacial geomorphic research has been conducted in the Uinta Mountains region. Other than hypotheses put forth by early workers on the integration of the Green River across the uplift (Powell, 1876; Sears, 1924; Bradley, 1936), the middle–late Cenozoic record has been explored mostly by W.R. Hansen (e.g., Hansen, 1965, 1984, 1986). He had several keen and interesting observations about regional landscape evolution, though none of his work focused on the Quaternary. Osborn (1973) and Nelson and Osborn (1991) mapped and employed relative dating of soil profiles on the extensive series of outwash terraces on the south flank of the Uintas. They distinguished at least twelve outwash surfaces in the Uinta Basin, some hypothetically pre-dating the Quaternary, and estimated relatively high incision rates of 170–320 m/m.y. In terms of glacial geology, W.W. Atwood completed the first intensive study of the Uintas. In his seminal 1909 monograph, Glaciation of the Wasatch and Uinta Mountains, he described the glacial geomorphology of every major drainage in the range and identified evidence for two major glacial advances. Later, in a report on the geomorphology of the north slope, Bradley (1936) recognized three separate ice advances; from oldest to youngest, he named these the Little Dry, Blacks Fork, and Smiths Fork glaciations after the valleys where their deposits were particularly well preserved. In his summary report on glaciation of the Rocky Mountains, Richmond (1965) correlated the Little Dry deposits with the Illinoian and pre-Illinoian glacial stages of the U.S. Midwest, the Blacks Fork with the early Wisconsin, and the Smiths Fork with the late Wisconsin glacial stage. He also correlated the Blacks Fork advance with the Bull Lake glaciation in the Wind River Range and the Smiths Fork glaciation with the Pinedale. More recent mapping has also considered the Blacks Fork glaciation to be correlative with the Bull Lake glaciation (Richmond, 1986; Bryant, 1992), although the Bull Lake glaciation has been shown to pre-date the early Wisconsin (Chadwick et al., 1997; Sharp et al., 2003). Thus, the overall glacial frame-

New Quaternary research in the Uinta Mountains work in the Uintas mirrors that of numerous other western ranges with a well-preserved MIS-2 advance (Smiths Fork glaciation), and a slightly more extensive older advance, considered to have occurred during MIS-6 (Blacks Fork glaciation). In the past few decades, several unpublished theses have focused on details of the glacial geology and geomorphology in specific areas of the Uintas (e.g., Schoenfeld, 1969; Barnhardt, 1973; Osborn, 1973; Grogger, 1974; Gilmer, 1986; Schlenker, 1988; Zimmer, 1996; Douglass, 2000). Most recently, Munroe (2001) mapped the glacial geomorphology of the northern Uintas at 1:24,000 scale and reconstructed the late Quaternary glacial history of this area on the basis of glacial and periglacial landforms. Similarly, Laabs (2004) mapped the south slope at 1:24,000 scale and applied cosmogenic surface exposure dating and numerical modeling of past glaciers in an attempt to better constrain the timing and paleoclimate of the local LGM. On the basis of thirteen cosmogenic 10Be exposure ages of boulders on adjacent moraines, Laabs (2004) determined that the Smiths Fork glacial maximum ended in the south-central Uinta Mountains by 17.6 ± 1.1 ka. Recent investigations into the fluvial record in the Uinta Mountains are fewer, and have largely concentrated on sediment transport and historic channel adjustments to dams, water diversions, and other human impacts. Ringen (1984) and Lenfest and Ringen (1985) quantified suspended sediment-discharge relationships for a number of stream gages on tributaries. Brink and Schmidt (1996) evaluated the historic record of geomorphic change on the Duchesne River, and a series of subsequent studies focused on that drainage. This includes studies of bedload transport, channel form, and channel migration (Smelser, 1997; Paepke, 2001), and the geomorphic effects of dams and water diversion in terms of sediment storage and channel response (Stamp, 2000; Gaeuman et al., 2005). The historic fluvial geomorphology of the Green River in the eastern Uinta Mountains has been investigated by Andrews (1986) and Grams and Schmidt (1999), especially in terms of sediment budget and the influence of Flaming Gorge dam. Finally, Martin (2000) and Larsen et al. (2004) studied debris-flow activity and the effects of low-magnitude floods on debris fans in the Green River canyons of Dinosaur National Monument. DAY 1—SALT LAKE CITY TO VERNAL: THE SOUTHERN UINTA MOUNTAINS Introduction The first day of this field trip (total driving distance ≈247 mi [396 km]) focuses on the glacial geology of the southern Uinta Mountains. The trip crosses the Wasatch Mountains and proceeds eastward across the south flank of the Uintas, ending in Vernal, Utah. Stops in the Lake Fork and Yellowstone River drainages are to observe Pleistocene glacial deposits and landforms along with a few prominent mass-wasting features. Two stops include moderate hiking and climbing at elevations up to 3323 m above

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sea level (m asl; 10,900 ft asl). Please be prepared for hiking in rough and possibly wet terrain and for sudden changes in weather (temperatures at high elevations may be close to 0 °C). Access to the last stop (Stop 1.6, Lake Fork Mountain) may be prohibited by inclement weather or adverse road conditions. Directions to Stop 1.1 From the Salt Palace Convention Center in downtown Salt Lake City, drive west on E 400 St./Utah Hwy 186 to South Main St. After 0.1 mi, turn left to follow Utah Hwy 186. After 0.1 mi, turn right onto South State St./U.S. Hwy 89. After 2.9 mi, merge onto I-80. Enter Parley’s Canyon. After traveling 23.1 mi on I-80, merge onto U.S. Hwy 40 via Exit 148 toward Heber and Vernal and proceed 70 mi. In Duchesne, turn left onto Utah Hwy 87 East and zero your trip odometer to begin the road log for Day 1. At mile 15.5, turn left onto road 21000 W and follow signs for Mountain Home, Moon Lake, and Yellowstone Canyon. Just prior to mile 17.4, the road descends to an outwash surface mapped by Nelson and Osborn (1991) as pre–Bull Lake in age. Turn right at mile 19.1 onto an unnamed county road toward Moon Lake. This road enters Ute tribal land at mile 19.8; be advised that stopping vehicles or exiting the road are prohibited in this area. At mile 20.5, the road ascends to an Altonah-age outwash surface, descends through a landslide in this outwash at mile 21.5, and descends to a Smiths Fork–age outwash at mile 23. Here, the scarp on the left side is cut into pre–Smiths Fork outwash. After mile 24.8, the road ascends to a Smiths Fork–age outwash fan (which is pitted near its head) and crosses Smiths Fork–age moraines at miles 26.2 and 27. At mile 28.3, park along the right side of the road just north of the Ashley National Forest boundary. Hike 0.4 mi due east to the top of the small hill to observe moraines, bedrock structures, and mass wasting features near the mouth of Lake Fork Canyon. Stop 1.1—Lake Fork Canyon (Mile 28.3) This stop provides a spectacular view of the moraine sequence at the mouths of Lake Fork and Yellowstone River canyons. The small bedrock-cored hill we have climbed is composed of Triassic-Jurassic sedimentary rocks of the Ankareh Formation and Nugget Sandstone (Bryant, 1992). You likely noticed sandstone and orthoquartzite erratics of the Uinta Mountain Group during the ascent; these are derived from Blacks Fork–age till that caps most of the hill. Although some of the erratics may have been deposited during the last glaciation, this hill likely split the termini of the Lake Fork glacier. This has led to excellent preservation of moraines on its west side (where the terminal moraine has not been subject to erosion by the Lake Fork River). Looking to the west and south across the Twin Pots Reservoir affords views of the moraine sequence at the mouth of Lake Fork Canyon (Figs. 2 and 3). The low-relief (less than 10 m) moraine immediately south of Twin Pots Reservoir is a Smiths Fork recessional moraine; the higher relief (more than 70 m), Ponderosa pine–covered, compound ridge beyond it is the prominent, sharpcrested Smiths Fork terminal moraine. Nine cosmogenic 10Be

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ages of this moraine from Laabs (2004) indicate this moraine was deposited prior to 16.9 ± 0.4 ka (based on a production rate of 5.1 atoms g SiO2−1 yr−1 adjusted for altitude and latitude by using scaling factors from Stone [2000]), which approximates the termination of the Smiths Fork glacial maximum in this valley. This moraine is distinguished from the outer Blacks Fork moraine (which is not visible at this stop) by its sharp crests and steep distal slopes, higher frequency of moraine surface boulders, and the presence of an ice marginal drainage valley that exceeds 60 m depth in places (Fig. 3; see also Laabs and Carson, 2005). The Smiths Fork moraine grades to an outwash fan that is more than 20 m below a fan deposited during the Blacks Fork glaciation (Fig. 3). The outermost moraines visible to the south likely correlate to the Altonah moraine; these high-relief moraine ridges grade into outwash terraces that are more than 50 m above mod-

ern grade (Laabs and Carson, 2005). The age of this episode is poorly known, but may correlate to the Sacagawea Ridge glaciation believed to be correlative to marine oxygen-isotope stage 16 (Chadwick et al., 1997). Outcrops to the east of the Lake Fork River expose an unconformity between south-dipping Mesozoic sedimentary rocks (namely the Triassic Woodside, Thaynes, and Ankareh Formations; the Triassic-Jurassic Nugget Sandstone; and the Jurassic Preuss Sandstone and Twin Creeks Limestone) and Oligocene sandstones and conglomerates of the Starr Flat Member of the Duchesne River Formation (Bryant, 1992). The Starr Flat Member was deposited by ancient tributaries of the Duchesne River while the Uinta Mountains were rising during late stages of the Laramide Orogeny. This unit may be correlative in places near Lake Fork canyon to the Bishop Conglomerate, a tuffaceous

Figure 2. Map of field trip stops on Day 1 (black dots). Moraines deposited during the Smiths Fork (solid black lines), Blacks Fork (short-dashed lines), and Altonah glaciations (long-dashed lines) are shown (from Laabs, 2004). Notable locations are Fish Creek (FC), Raspberry Draw (RD), Harmston Basin (HB), and Cow Canyon (CC). Shaded topography is from portions of the U.S. Geological Survey 15′ Kings Peak and Duchesne quadrangles.

New Quaternary research in the Uinta Mountains sandstone and conglomerate that overlies the Gilbert Peak erosion surface (which will be discussed at Stop 2.1). Views to the north display the broad U-shape of the Lake Fork canyon. The western side of the canyon is covered mainly by till immediately north of this stop. Lateral moraines high above the valley bottom are easily visible and represent deposition during the Blacks Fork and Smiths Fork glaciations. The eastern side of the valley contains little till near the valley mouth, perhaps due to oversteepening of the valley side and subsequent mass wasting caused by toe-slope erosion by the Lake Fork River. Post-glacial mass wasting has taken place along both sides of the valley, as indicated by abundant rock falls, slumps, and landslides on the east side of the valley and dissected moraines on the west side.

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Directions to Stop 1.2 Continue driving northwest toward Moon Lake. At the bridge crossing mile 28.5, notice at 3 o’clock the massive slope failure between two limestone ridges on the east wall of Lake Fork Canyon. Landslides are common on this side of the valley, which has been oversteepened near the mouth of the canyon due to toe-slope erosion by the Lake Fork River or by glacial erosion. After the bridge crossing, till covers most of the valley to the left of the road until mile 30, where the road ascends upon the broad alluvial fan discussed at Stop 1.1 (Fig. 2). At mile 31, a large landslide in Raspberry Draw is viewable to the right. Stop 1.2 is at the sign that indicates Raspberry Draw at mile 31.4; park along the right side of the road.

Figure 3. (A) Map of glacial deposits at the mouth of Lake Fork canyon (from Laabs and Carson, 2005). (B) View of moraine sequence to the south of Stop 1.1. A—Altonah moraine; SF—Smiths Fork moraine; SFR—Smiths Fork recessional moraine; TP—Twin Pots reservoir.

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Stop 1.2—Raspberry Draw (Mile 31.4) A large landslide in Raspberry Draw, an eastern tributary of the Lake Fork River (see Fig. 2 for location), cuts across two lateral moraines on the side of Lake Fork Canyon. The moraines are discontinuous and their ages are poorly known, but their crests are at elevations very close to Blacks Fork and Smiths Fork lateral moraines on the other side of the valley (Laabs, 2004); this indicates that the landslide occurred sometime after ca. 17 ka. The surface of the landslide is hummocky and several sag ponds are present, but it is heavily grazed by cattle that have likely disturbed pond sediments that might otherwise provide a limiting radiocarbon age of this feature. The origin of the landslide is undoubtedly related to bedrock geology; the underlying strata are limestones and dolomites of the Mississippian Humbug Formation and shales of the Mississippian Doughnut Formation. As mentioned, landslides are common at the contact between these two units and are most prominent in Raspberry Draw and in Cow Canyon; the latter is a western tributary of the Yellowstone River (see Fig. 2 for locations; Laabs, 2004). One possible explanation for the origin of these landslides is saturation of weakly-cemented, fractured shales of the Doughnut Formation where they overlie low-permeability dolomite and/or limestone of the Humbug Formation. Increased pore-water pressure in saturated material weakens its shear strength and can cause it to move down slope, perhaps in response to oversteepening of the slope or seismic shaking (seismic shaking alone can also reduce the sheer strength of saturated material). Another possibility is that the lateral moraines north of the valley were once continuous across Raspberry Draw and dammed a small lake in the tributary valley. Failure of the moraine dam, the lake-bottom sediments, and underlying lowstrength rocks may have triggered the landslide. Small, morainedammed lakes are common in tributaries to the Lake Fork and Yellowstone Valleys; however, distal slopes of the remaining lateral moraines on the north side of Raspberry Draw reveal no lake sediment. Directions to Stop 1.3 Continue driving northwest across the alluvial fan toward Moon Lake. At mile 32.2, the head of the broad alluvial fan can be seen to the left. Flat-lying sandstone and conglomerates of the Duchesne River Formation seen above the head of the fan are the primary source of the fan sediment. Proceed northward to mile 32.2, turn right, and drive to the end of the driveway; this is Stop 1.3. Stop 1.3—Lake Fork River Cutbank below Moon Lake Dam (Mile 32.3) Moon Lake is a reservoir that enlarged a natural lake when an earthen dam was constructed just south of the original outflow in 1938. The lake maintains its original shape of a crescent moon (Fig. 2) but is now almost 6 km long (compared to an original length of ~5 km; Atwood, 1909) and has an average depth of ~14 m. Outflow from the dam generates hydroelectri-

city utilized locally, and discharge is used primarily for irrigation in the Uinta Basin. The origin of Moon Lake is likely related to deposition of the alluvial fan on the western side of the valley (Fig. 2) and a back-rotational slump in bedrock on the east side, both of which are visible here. We are standing on the distal edge of the alluvial fan, which is composed primarily of debris-flow diamicton exposed in a cutbank just below this site (note: the trails that descend upon the cutbank are highly unstable; it is strongly recommended that you view it from above and stay off the trails). The source of this sediment is sandstone and conglomerate of the Duchesne River Formation exposed at the top of the high ridge to the west. Physical properties of the alluvial fan sediment strongly resemble those of its source rock, which contains medium- to coarse-grained sand and abundant rounded cobbles and boulders and is actively eroding today. Based on an observed breach in the Smiths Fork right-lateral moraine to the west and a lack of buried soils, erosional contacts, and glacial till in the cutbank, we believe that all of the fan sediment exposed here was deposited after the last glaciation. The back-rotational slump across the valley from this stop is in Madison Limestone, Tintic Formation, and Red Pine Shale; mass wasting in the latter unit is ubiquitous in the southern Uinta Mountains. Evidence for back rotation of the slump can be viewed to the east where beds of the Tintic Formation and Madison Limestone are dipping southward more steeply than in situ strata. The slump extends westward into Moon Lake, and part of it was probably excavated during construction of the dam. The dam is therefore built into alluvial fan sediment on the east side of the valley and a slump deposit on the west side of the valley. Atwood (1909) observed Moon Lake in its original form and did not describe evidence of a moraine dam or bedrock sill that could have formed a threshold for Moon Lake; therefore, it is not considered a paternoster lake. Northeast of Moon Lake dam, prominent lateral moraines are visible on the sides of Fish Creek Valley, an eastern tributary of the Lake Fork River. These were deposited during the Smiths Fork glaciation when a small tributary glacier advanced southward from a broad high basin and intersected the main valley glacier. The headwaters of Fish Creek, a former glacier accumulation area, will be viewed from Stop 1.6. In summary, the Quaternary history of lower Lake Fork Canyon includes repeated glaciations during the middle and late Pleistocene, the last of these terminated by ca. 17 ka (Laabs, 2004). Post-glacial events include mass wasting on steep valley slopes, erosion and mass wasting of lateral moraines, and deposition of the alluvial fan on the west side of the canyon. The Lake Fork River has incised its valley since the Smiths Fork glaciation, leaving an outwash terrace ~8 m above modern grade. Directions to Stop 1.4 Return to the paved road, turn right, and enter the Moon Lake campground at mile 33.6. Restrooms and water are available at this lunch spot. When exiting the campground, reset your odometer to zero at the cattle guard for accurate mileage to afternoon

New Quaternary research in the Uinta Mountains stops. Drive southeast on the unnamed county road past all of the previous stops and leave Ashley National Forest. Turn left at mile 10.1 and drive east on the bridge over Lake Fork River. Smiths Fork–age outwash gravels appear in roadcuts at mile 10.4. Turn left at mile 10.5 and drive north toward the Yellowstone campgrounds in Ashley National Forest. Looking west at mile 11.0 provides a cross-sectional view of the moraine sequence at the mouth of Lake Fork Canyon (described at Stop 1.1). Immediately east of this road is a Blacks Fork–age latero-frontal moraine at the mouth of Yellowstone Canyon; parts of this moraine were deposited on Tertiary strata and sediment exposed in places immediately east of the road. At mile 12.5, the relatively deep gorge (up to 35 m) to the left was an ice-marginal drainage valley occupied during the last glaciation; this valley can be traced northward to the Ashley National Forest boundary. Enter Ashley National Forest at mile 14.4 and bear right at the fork at mile 14.6; here, the road descends onto Smiths Fork–age recessional moraines and kame terraces deposited during the last deglaciation. Stop 1.4 is the pullout area on the right at mile 15.3. Stop 1.4—Moraine Sequence in Yellowstone Canyon (Mile 15.3) Latero-frontal moraines that represent glacier termini during three Quaternary glaciations are well preserved and easily viewable just south of this stop. Lateral moraines on the east side of the valley are best preserved and are separated by broad ice-marginal drainage valleys. Smiths Fork terminal moraines are visible on both sides of the valley, where they locally have over 70 m of relief (Figs. 2 and 4). A well preserved sequence of Smiths Fork recessional moraines covers much of the valley center between terminal moraines; some of these features are clearly visible to the east. Part of a Blacks Fork moraine can be seen down valley to the south (Fig. 4), where a latero-frontal moraine curves toward the center of the valley. The outermost moraine is visible to the southeast and is the type section for the Altonah glaciation, where a prominent left-lateral moraine rises ~90 m above Mud Spring Draw (Fig. 4). Laabs et al. (2004) attained seven cosmogenic 10Be exposure ages of moraine boulders on the Smiths Fork left-lateral moraine, which is the highest visible ridge due east of this stop. The weighted-mean exposure age of four of these boulders (three samples were excluded based on statistics and boulder height) is 18.0 ± 0.7 ka, which they interpret to represent the end of the Smiths Fork maximum in this valley. This age estimate overlaps the weighted mean age of the Smiths Fork terminal moraine in the Lake Fork Canyon (16.9 ± 0.4 ka), and several boulder-exposure ages from these two moraines overlap within 2σ uncertainty. Laabs (2004) computed a weighted mean age of 13 samples from both moraines of 17.6 ± 1.1 ka, which represents the termination of the Smiths Fork maximum in the southwestern Uintas. The Yellowstone River flows east of, and more than 50 m below, this stop. Glacial sediment has filled in much of the west side of Yellowstone Canyon; this stop provides a view of recessional moraines that comprise the surface deposits between

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Yellowstone River and the Smiths Fork latero-frontal moraine (Figs. 2, 4, and 5). Westlund (2005) mapped nine recessional moraines in this area, some of which are breached by Crystal Creek on the west side of the valley. This moraine sequence represents deposition during the last deglaciation; the Smiths Fork terminal moraine is visible upslope to the west (it is the ridge covered with prominent Ponderosa pines). Some recessional moraines are separated by kame terraces (Fig. 5), one of which is exposed in a road cut just east of the parking area. Here, glacial sediments are distinguished from tills elsewhere in the valley by the grain-size distribution of the matrix sediment (sandy silt versus silty sand) and by the presence of carbonate in the matrix and as coatings on cobbles and boulders. Clastic carbonate grains are rare in tills of the Uintas, and secondary carbonate is common only in pre–Smiths Fork tills. The source of carbonate in the kame terrace sediments is outcrops of Madison Limestone near the head of Hells Canyon (see Fig. 2 for location). Indeed, this implies that Hells Canyon was the source of ice-marginal drainage sediment on the west side of Yellowstone Canyon during the last glaciation; the evolution of fluvial drainage in this area will be discussed at the next stop. Directions to Stop 1.5 Turn around and drive south-southeast to the fork encountered earlier. At mile 15.9, turn right (the turn is actually ~300° to the right) and follow FR 227 toward Hells Canyon. At mile 16.2, the Smiths Fork–age right-lateral moraine (covered by Ponderosa pines) can be viewed above the parking area to the west. This ridge, along with multiple recessional moraines below it, is cross-cut by Crystal Creek at mile 16.5; after crossing Crystal Creek, the Smiths Fork–age moraines are to the east of the road. At mile 16.6, outcrops of colluvium and landslide diamicton on top of a buried soil in Blacks Fork till can be viewed on the left side of the road. The road enters an ice-marginal drainage valley at mile 16.7 that separates the Smiths Fork moraines on the right from a Blacks Fork moraine on the left. At mile 18.0, park in the campsite on the right side of the road and climb on to the ridge on the east side of the campsite; this is Stop 1.5. Stop 1.5—Ice-Marginal Drainage in Yellowstone Canyon (Mile 18.0) The stream in Hells Canyon is visible directly north of this stop where it flows eastward to the Yellowstone River. The presence of an ice-marginal drainage valley to the south, however, suggests that drainage was to the south during the last glaciation and that the lateral moraine here once connected to the ridge on the north side of Hells Canyon (Fig. 6). A continuous lateral moraine would have blocked eastward drainage out of Hells Canyon and diverted flow southward along the distal side of the moraine, as evidenced by the presence of sand and rounded cobbles on the bottom of the valley through which we have been driving. Southward flow extended at least 3 mi south of this stop where the stream incised a relatively deep valley that was seen on the drive to Stop 1.4. Depending on the surface elevation of the

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Figure 4. (A) Map of glacial deposits at the mouth of Yellowstone canyon. (B) View of moraine sequence to the south of Stop 1.4. This photo was taken at the top of a moraine just west of Stop 1.4). BF—Blacks Fork moraine; SF—Smiths Fork moraine; SFR—Smiths Fork recessional moraine. From Laabs and Carson (2005).

lateral moraine that has been removed from this area, drainage from Hells Canyon may have been ponded by the moraine and back-filled northward into Harmston Basin (see Fig. 2 for location; see also Fig. 6). Therefore, topographically closed depressions in Harmston Basin may contain a record of drainage from Hells Canyon during the last glaciation. The stream in Hells Canyon breached the Smiths Fork lateral moraine here during the last deglaciation, flowed southward along the next successive lateral moraine to deposit kame terrace sediments, and then breached each successive lateral moraine

and flowed southward before cutting through the youngest recessional moraine and adopting its current path eastward to Yellowstone River. As noted above, the presence of carbonate in kame terrace sediments indicates drainage from Hells Canyon along with the fact that no other source of ice-marginal drainage was available to deposit these sediments. Directions to Stop 1.6 Continue driving along FR 227 and enter Hells Canyon at mile 18.6. The stream in Hells Canyon is ephemeral in the lower

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top of Lake Fork Mountain (note: this hike ascends 400 vertical feet [122 m] at high elevations; although the climb is not steep, please use caution if you are not accustomed to hiking at elevations above 3000 m).

Figure 5. Cross-sectional view of Smiths Fork lateral moraines and kame terraces west of Stop 1.4. This photo was taken at the top of a moraine just west of Stop 1.4. SF—Smiths Fork lateral moraine; SFR—Smiths Fork recessional moraine; SFK—Smiths Fork kame terrace.

part of the canyon, partly because some of the drainage is lost to the subsurface through sinkholes in the Madison Limestone. Note the abundance of sandstone and orthoquartzite cobbles and boulders in the talus along the right side of the road; these are derived from conglomerate and diamictite of the Duchesne River Formation. At mile 23.5, turn left onto FR 196 toward Mill Park. Weathered colluvial deposits derived from local outcrops of Tertiary gravels can be seen on either side of the road and at Mill Park. Lake Fork Mountain can be seen to the west across Mill Park at the junction at mile 24.4. Turn left at the junction, proceed ~50 ft, and turn right onto an unnamed forest road. The road intersects an ATV trail at mile 25.4; parking here is recommended, although the road continues toward the top of Lake Fork Mountain and can be driven in a high-clearance vehicle. If the road is wet, a four-wheel drive vehicle may be needed to continue driving past this point. Park near the intersection with the ATV trail and hike ~1 mi along the road to an air quality station at the

Stop 1.6—Lake Fork Mountain (Mile 25.4) This site provides a broad view of the geomorphology of the southern Uintas and a rare opportunity to view terminal moraines and cirque headwalls from a single location that is (almost) accessible by car (Fig. 7). Lake Fork Mountain is an east-west–trending strike ridge of Madison Limestone bounded to the north by a normal fault. The Madison Limestone rarely forms continuous hogback ridges due to gently dipping bedrock in the southern Uintas, which contrasts with steeply dipping strata in the northern Uintas where hogbacks are common (see Stops 3.1 and 3.4). Looking southward from Lake Fork Mountain affords views of the aforementioned moraine and outwash sequences on the piedmont below Lake Fork and Yellowstone River canyons. On a clear day, one can see across the Uinta Basin to the uplands of the San Rafael Swell. High peaks, cirques, and broad valleys in the headwaters of Lake Fork and Yellowstone River canyons can be viewed to the northwest and northeast, respectively (Fig. 7). Broad, compound cirques that exceed 10 km in width (e.g., Brown Duck and Yellowstone basins; Fig. 7) and ultimately drain into either Lake Fork or Yellowstone River canyons are clearly visible to the northwest and northeast, respectively, and are representative of the distinct morphology of paleo-glacier accumulation zones in the southern Uintas. Such valleys must have contained many small glaciers at the onset of past glaciations that ultimately coalesced and joined the main valley glacier. Moraines in these valleys only exist in small cirques, and glacial deposits vary from several meters of till to a scattering of boulders over bedrock. The morphology of these high basins is distinct from those on the north slope (as viewed at Stop 3.1), where valleys are narrower and ice masses coalesced into smaller individual glaciers during the Smiths Fork glaciation. Although high basins lack topographic shading and receive higher amounts of solar

Figure 6. Photo of Harmston Basin as viewed to the north of Stop 1.6. BF— Blacks Fork–age moraine; SF—Smiths Fork–age moraine; SFR—Smiths Fork recessional moraine.

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Figure 7. Panoramic photo of the high Uintas as seen from the top of Lake Fork Mountain. Top image shows the headwaters of Lake Fork basin as viewed to the northwest; lower image shows the headwaters of Yellowstone River basin as viewed to the northeast.

radiation than valleys on the north side, they are at elevations above the reconstructed equilibrium line altitude of the Smiths Fork maximum (~3000 m; Shakun, 2003; Laabs and Carson, 2005). The northeastern part of the Yellowstone drainage basin roughly coincides with the domal center of the Uinta Mountain anticline (see Stop 2.1, next section), which could explain why this is one of the highest areas in the Uinta Mountains and could therefore sustain ice masses in broad, poorly shaded basins under glacial climates. Directions to Vernal Descend Lake Fork Mountain and proceed southward. Do not turn right at mile 39.0; instead, proceed south toward Altonah. The Piñon-covered hills to the east at mile 39.9 are part of the Blacks Fork–age terminal moraine. The road climbs onto Smiths Fork–age outwash at mile 41.1 then ascends to Blacks Fork–age outwash at mile 41.5. At mile 42.2, the Blacks Fork–age terminal moraine is the tree-covered surface that rises above the outwash to the north. The Altonah left-lateral moraine comes into view at 10 o’clock just before mile 43.6. At mile 46.0, turn left at the stop sign onto 7000 North. Turn right onto 16000 West at mile 47.0 and proceed southward to a stop sign at mile 50.0. Turn left onto Utah Hwy 87 East, drive through Altamont, and bear right at mile 50.8 toward the towns of Mount Emmons and Upalco. Proceed on Hwy 87 and turn left onto U.S. Hwy 40 at mile 67.0. Bear right at mile 72.5 and follow signs to Vernal.

DAY 2—VERNAL TO RED CANYON LODGE—THE EASTERN UINTA MOUNTAINS Introduction Day 2 (total driving distance ≈106 mi [171 km]) focuses on the nonglacial Tertiary-Quaternary landscape evolution of the eastern Uinta Mountains. Ideas on the classic enigma of the integration of the Green River across the eastern Uinta uplift are reviewed, and new surficial mapping in western Browns Park is presented, which features deposits interpreted as a middle Pleistocene paleoflood and impoundment deposit associated with a landslide dam across the paleo–Green River. Directions to Stop 2.1 Start in Vernal at the intersection of Main Street and 500 North (mile = 0.0). Turn east on 500 North, which becomes Jones Hole Road. Take the left fork after crossing Ashley Creek, continue over the south flank of the Buckskin Hills, and stay left at the intersection with Brush Creek Road at mile 5.7. At mile 6.5, note the unstudied sequence of Plio(?)-Pleistocene gravels and terraces along Brush Creek and at the top of the Buckskin Hills. Nelson and Osborn’s (1991) work on the south flank to the west of here estimated relatively high Quaternary incision rates using soil-development indices on alluvial outwash surfaces. Stay right to cross Brush Creek, and climb to the Diamond Mountain

New Quaternary research in the Uinta Mountains Plateau. At mile 13.8, the road has been damaged by landsliding of the Cretaceous Mancos shale. Visible at this point across the piedmont to the southeast, the Green River has incised the Split Mountain Anticline. The top of this Laramide structure is beveled by the Gilbert Peak erosion surface and capped by the Bishop Conglomerate, which is what the paleoriver probably flowed upon, enabling superimposition of the drainage. Stop 2.1 (Mile 15.0) This roadcut is an exposure of a tuffaceous facies in the middle of the Bishop Conglomerate atop the Diamond Mountain Plateau. Concordant plateaus and summits all around record a low relief post-Laramide landscape (Fig. 8). Two Tertiary deposits, the Bishop Conglomerate with its underlying Gilbert Peak erosion surface (seen here) and the Browns Park Formation (seen at next stop), provide a framework for understanding post-Laramide structural deformation, drainage development, and subsequent erosion of the eastern Uintas. The Gilbert Peak erosion surface is an extensive pediment that formed on the north, east, and south flanks of the range during a prolonged period of tectonic stability in Oligocene time. The Bishop Conglomerate was deposited basinward and atop the pediment as erosion continued at the retreating mountain front. An ash at this locality has produced a new Ar/Ar date of 30.54 ± 0.22 Ma (Kowallis et al., 2005). Evidence for the regional tilting of the Bishop Conglomerate–Gilbert Peak surface lead Hansen (1986) to conclude that the northeastern part of the Uinta fault was reactivated with normal dip-slip motion in Neogene time. In Hansen’s model, this and faulting associated with the formation of the Browns Park graben represent the gravitational collapse of the eastern part of the Uinta uplift (Bradley, 1936; Hansen, 1984, 1986). Note that the surface of the Bishop Conglomerate here on the Diamond Mountain Plateau has multiple terrace levels (Fig. 8). There has been significant later Tertiary and Quaternary reworking and beveling of this inherited planation surface.

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Additional evidence for late Cenozoic deformation in the eastern Uinta Mountains includes the courses of numerous streams that have been altered and reversed (Hansen, 1984, 1986). Consequent streams flowed southward through the Diamond Mountain Plateau from the crest of the Laramide uplift in what is today the area of Browns Park. The courses of many of these drainages have now been reversed, as Hansen (1986) pointed out through evidence of an older, relict drainage divide and the barbed pattern of southerly flowing tributaries that make nearly 180° turns before joining northward-flowing trunk streams. The tilting and subsidence of Browns Park developed steep drainages that have eroded headward and captured southflowing drainages, thus displacing the drainage divide southward (Fig. 9). This process continues today, as apparent along the upcoming field trip path, where Crouse Creek and Jackson Draw are very close to capturing the headwaters of Pot Creek. Directions to Stop 2.2 Continue north and then east on paved Jones Hole Road across the Diamond Mountain Plateau. At mile 22.6, turn left onto gravel Crouse Canyon/Browns Park Road. Note the gently dipping hogback ridges held up by the Pennsylvanian Round Valley Limestone and the Mississippian Madison Limestone as we cross the beveled and buried southern flank of the Uinta uplift. Near the top of a Madison hogback, at mile 24.8, stay left at the fork (Crouse Canyon Road) and cross a drainage divide/windgap capped by Bishop Conglomerate(?). Ahead, to the north, broad valleys cut in UMG sandstones are filled with what has been mapped as Bishop Conglomerate. This is an example of the relict Eocene-Oligocene drainages Hansen (1986) interpreted as once flowing to the south (e.g., through this pass) from the high Laramide range crest, and which then foundered with the range, driving drainage change. At mile 27.6, turn right on the intersecting gravel road and head east to mile 29.9, where we take a left turn and go north at

Figure 8. Photographic panel looking west from the south flank of Diamond Mountain Plateau toward the location of Stop 2.1. The broad topographic shoulder of the range is a bedrock pediment near the axis of the uplift, but it becomes progressively overlain by Oligocene gravel toward the basin margin, as epitomized by the Diamond Mountain Plateau.

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the T. At mile 31.5, cross a subtle drainage divide and descend into the incised head of the Crouse Creek drainage, which is poised to capture the upper Pot Creek drainage system behind us. Note the last outcrops of mapped Bishop Conglomerate at mile 37.0, as we descend into the steepest reach of the youthful Crouse Canyon and out of the broad, relict Oligocene valley. Pit Draw, to the east, is an example of a barbed drainage meeting Crouse Creek. At mile 38.3, note significant talus with possible protalus ramparts (?) on the east side of the canyon. Turn left at mile 41.1 onto Taylors Flat Road after exiting the canyon mouth, cross Crouse Creek, and climb past roadcut exposures of tuffaceous

Figure 9. Map illustration of Green River integration history. (A) Oligocene-Miocene landscape with location of original paleodrainage divide indicated by dashed line. After Miocene subsidence of Browns Park and collapse of the eastern uplift, the upper Green River was diverted into the graben, and a tributary drainage of the paleo-Yampa underwent subsidence of its headwaters. (B) In post–Browns Park time (Pliocene?), the Green River was captured or diverted through the Canyon of Lodore, and the drainage divide shifted south to its present position in the eastern Uintas.

Browns Park Formation. Continue to mile 42.7, a roadside stop overlooking Swallow Canyon. Stop 2.2 (Mile 42.7) This stop provides an overview of where the Green River passes through a series of geologic transitions. Just west of Browns Park, the steep walls of Red Canyon temporarily open into a terraced, park-like area encompassing Little Hole and Devils Hole (Fig. 9). Farther downstream, the valley becomes narrow again as the Green River continues through the well-indurated UMG sandstones of lower Red Canyon. The canyon opens into western Browns Park, where bedrock changes to the weaklyconsolidated tuffaceous sediment of the Tertiary Browns Park Formation. The Green River passes through the short bedrock gorge of Swallow Canyon at this stop and then continues through the heart of Browns Park. Quaternary landforms and deposits in Browns Park show evidence of minor displacement along the Home Mountain fault at the linear northwest basin edge, whereas the southern mountain front is generally highly sinuous and embayed. About 12 mi downstream at the center of Browns Park, the Green River turns south and punches through the Uinta uplift at the Gates of Lodore (Fig. 9). It then continues across the range past its confluence with the Yampa River, forming the deep canyons of Dinosaur Canyon National Monument. There is a long history of debate on the timing and mechanisms of integration of the Green River across the Uintas and into the greater Colorado River. It was here that John Wesley Powell coined the term “antecedent” to relate his hypothesis about how the Green River came to cross the Uinta uplift and cut the Canyon of Lodore. This idea that the river path is older than a relatively young and active orogen was subsequently discarded in favor of some combination of superposition and stream capture. Swallow Canyon, seen here, is certainly an example of river superposition. Subsequent workers provided two hypotheses for the mechanisms of drainage integration across the Uinta uplift: (1) headward erosion of a small stream through the uplift and capture of the upper Green River drainage (Bradley, 1936; Hansen, 1965); and (2) superposition of the Green River over the eastern Uintas as it filled Browns Park with sediment and spilled over the range to the south (Sears, 1924; Hansen, 1986). These hypotheses each have problems. For example, headward erosion of a first-order stream through resistant sandstone is difficult, and there is no known evidence in the Browns Park Formation for sediment of a proto–Green River. A review and new perspective on these options provides a preferred hypothesis for drainage integration (Pederson and Hadder, 2005). Hansen (1969) recognized that the subsidence associated with Miocene extensional collapse of the eastern Uintas must have played a key role in capturing or diverting an upper Green River south into Browns Park, as well as south from Browns Park across the eastern Uinta uplift. The paleo–Green River, once diverted into the subsiding Browns Park, would have found a gentle path continuing along the strike of the graben into northwestern Colorado (Fig. 9). But the path of the Canyon

New Quaternary research in the Uinta Mountains of Lodore follows the trend and spacing of the relict, subsided drainage headwaters of the Diamond Mountain Plateau, such as the one just followed along Crouse Canyon. The Laramide (pre–Browns Park) drainage divide would have extended into the heart of Browns Park in the area of the Gates of Lodore (Bradley, 1936; Hansen, 1986) (Fig. 9). That is, relict Oligocene topography suggests that the headwater drainage divides of a steeper, south-flowing tributary of the paleo-Yampa subsided into the Browns Park graben during Miocene faulting. This would have greatly eased the path for headward erosion of the tributary or for the spilling or diversion of the nascent upper Green River through the fallen divide atop the Browns Park Formation in late Miocene or Pliocene time (Fig. 9). The surficial stratigraphy here in western Browns Park is being studied, in part, with the goal of quantifying incision rates. Given the identification of the Lava Creek B tephra (see Stop 2.4) and mapping correlations, Counts (2005) estimated a Quaternary incision rate of 80–115 m/m.y. in western Browns Park. This is only about half that calculated in most other places in the region, including many likewise based on the Lava Creek B tephra (Table 1). This relatively low incision rate estimate on the north flank of the Uintas also contrasts with estimates made just across on the southern flank of the Uintas (Nelson and Osborn, 1991). Similarly, though not necessarily related, GIS-based calculations indicate the total late Cenozoic exhumation of the Green River Basin to the north is approximately half that of the Uinta Basin to the south of the Uinta Range (Pederson, 2004) (Fig. 10). An explanation is required for this lower total exhumation and, if initial results hold true with more geochronology, lower incision rates. There is no evidence in the Browns Park area for differential climate changes or significantly active subsidence or faulting in the Quaternary that could have reduced total exhumation and incision rates. To some degree, less exhumation to the north may be a function of the more recent integration of the Green River Basin into the overall system, as discussed above. But considering that incision is ultimately driven by baselevel

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fall, reduced exhumation to the north may also reflect buffering of the region from baselevel fall that has driven incision elsewhere. The upper Green River is isolated in the headwaters of the greater Colorado River system, far from its ultimate baselevel at the Gulf of California. In addition, the upper Green River is located above several steep canyon reaches (Fig. 10). Any baselevel fall downstream, whether caused by epeirogenesis or river integration, would diffuse and get absorbed or “hung up” in canyon reaches where stream power is concentrated to meet the resistance of harder bedrock or coarse sediment loading. Directions to Stop 2.3 Continue west on the Taylor Flats Road. Cross Sears Creek at mile 48.7, and then cross the terrace associated with Quaternary alluvial gravel 5 (Qag5) for the next few miles to the scattered dwellings of Taylors Flats at mile 51.5. This landform dominates western Browns Park, and most of the dissected piedmont fans and slopes in the area grade to this level. Continue north across the Green River at Bridge Hollow at mile 52.8, turn left toward the Jarvie Ranch Historic Site, and pull into parking area. Stop 2.3 (Mile 53.1) Hike up a short but steep slope to south of Jarvie Ranch, over Browns Park Formation, to a very coarse basal Qag5 deposit that grades up to typical Green River gravel and is finally capped with distinct piedmont sand and gravel. This outcrop provides an overview of the Quaternary stratigraphy of the Green River and an introduction to the Qag5 flood deposit. There are at least nine stratigraphically distinct and inset mainstem Green River deposits preserved in western Browns Park. These gravel and sand deposits generally range from 2 to 12 m thick and are associated with strath or thin fill terraces that have tread heights up to 227 m above modern grade (Fig. 11). Based on local piedmont equivalents as well as the landscape position of the latest Pleistocene (Smiths Fork) glacial outwash in the Henrys Fork field area (visited on Day 3), we interpret the

TABLE 1. CALCULATED INCISION RATES FOR REGIONAL RIVERS Rivers

Location

Incision rate

Age control

Yampa River

Northwestern Colorado

110–150 m/m.y.

Lava Creek B ash

Reheis et al., 1991

Bighorn River

Bighorn Basin, Wyoming

160 m/m.y.

Lava Creek B ash Field Camp ash

Palmquist, 1983

Wind River Basin, Wyoming

150 m/m.y.

Lava Creek B ash

Chadwick et al, 1997

NW Uinta Basin

170–320 m/m.y.*

Soil-development indices

Green River

NE Uinta Mountains, Browns Park

90–115 m/m.y.

Lava Creek B ash

Counts, 2005

Henrys Fork

NE Uinta Mountains, Manila, Utah

80–110 m/m.y.

Correlation to dated Wind River stratigraphy

Counts, 2005

Wind River Lake Fork, Yellowstone, Uinta

Study

Nelson and Osborn, 1991

*Reported Uinta Basin rates are for the mountain piedmont for comparison to same on the northern flank of range. These are higher than those reported (Nelson and Osborn, 1991) near the center of the Uinta basin.

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Figure 10. Schematic cross section along Green River, with river long-profile digitized from Trimble (1924). “Dinosaur canyons” are Lodore, Whirlpool, and Split Mountain canyons of Dinosaur National Monument. Uinta and Green River basins show basin-averaged late Cenozoic exhumation calculated following the methods of Pederson et al. (2002). The striped interval of Uinta uplift is a range of peak elevations between a higher profile across the peaks in the central-western part of range and a lower profile taken across the eastern Uintas. The signal of baselevel lowering that has been driving regional incision should diffuse and dampen as it propagates upstream and encounters steep canyon reaches.

presence of Qag1 Green River gravels to be mostly below grade, deposited during the LGM. The sandy upper alluvium and terraces of Qag1 visible in the floodplain are probably Holocene in age, and the youngest of these (historic) have been studied by Grams and Schmidt (1999). Qag1, Qag3, and Qag4 are distinct from this viewpoint, and higher gravels are visible, but the predominant surficial deposits in Browns Park and upstream in the lower reaches of Red Canyon are Qag5 sand and gravel (Fig. 11). These differ strongly in character between here and the next stop in lower Red Canyon at Little Hole, but they are correlative in terms of their position within the Quaternary stratigraphy and their interpreted genesis. The Qag5 here is very distinct from typical Green River gravels. It is underlain by a planar strath cut in the Browns Park Formation 40–50 m above grade, and the deposit thins strongly downstream from over 25 m thick at the mouth of Red Canyon to 10 m thick near Stop 2.2. The exposure of Qag5 near Jarvie Ranch is 23 m thick (Fig. 12), and Qag5 has a fan-shaped geometry that spreads across much of the valley floor in this area. The basal 18 m at this locality is a clast-supported boulder gravel with boulders that are exclusively subangular UMG sandstone up to 3.3 m in intermediate-axis diameter, with a mean diameter of 0.84 m (Counts, 2005). The framework openings are filled with a matrix of sand and rounded pebbles and cobbles of polymictic Green River provenance. Boulders are abruptly absent above 18 m, grading sharply to better sorted and rounded pebble-cobble gravel of the Green River. The caliber and dominance of UMG clasts in the base of the Qag5 deposit decreases strongly downstream, starting with boulders up to 5 m in intermediate diameter at the mouth of Red Canyon. The placement of these largest boulders at the apex of a fan-shaped deposit indicates an origin

Figure 11. (A) Photograph looking upstream (west) from Jarvie Ranch at Green River terrace gravels and mouth of lower Red Canyon. A ~50m-thick shale formation within the Uinta Mountain Group outcrops at the base of the south slope in lower Red Canyon on the outer bank of the undercutting river. (B) Composite surveyed cross section from the Jarvie Ranch area (modified from Counts, 2005). Qag5 deposits dominate the valley bottom and are anomalous in their sedimentology, thickness, and extent. Qag—Quaternary alluvial gravel of mainstem drainages; Qal—alluvium of active channels and floodplains.

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Figure 12. Overview of Qag5 deposit exposed immediately north of the historic Jarvie Ranch in western Browns Park; river flows from left to right. Inset is close up of the boulder gravel at the base of the deposit (from Counts, 2005). Qag—Quaternary alluvial gravel of mainstem drainages; Tbp—Tertiary Browns Park Formation.

from lower Red Canyon. Hansen (1965) interpreted these as a coarse facies of the underlying Browns Park Formation adjacent to steep paleocanyon walls, and there are, in fact, such deposits in the distinct Browns Park Formation nearby. Landslide scars and deposits cover much of the south wall of lower Red Canyon between the confluence of Red Creek and the mouth of Red Canyon (Fig. 12). In this reach of the canyon, a ~50-m-thick formation of Neoproterozoic shale within the UMG is exposed at river level (Dehler et al., 2006). This creates a situation in which the outer bank of the river impinges upon shale, undercutting a steep slope of sandstone that rises nearly 800 m above the river. The surface of the massive landslide deposit, according to field relations, predates at least Qag3, and it has been modified by fluvial process into a sloping terrace that sits at an elevation ~40 m above the top of the Qag5 mainstem deposit here at Jarvie Ranch. However, its elevation is concordant with the top of the anomalous Qag5 deposit at Little Hole (next stop). We hypothesize this older landslide impounded the Green River and that the failure of this landslide dam is recorded in the basal Qag5 here in western Browns Park. Directions to Stop 2.4 Drive out of Jarvie Ranch to the east, straight past Bridge Hollow, and turn left at the T-intersection toward the mouth of Jesse Ewing Canyon at mile 54.8. At mile 56.3, the Home Mountain fault bounding Browns Park is crossed and the road ascends the steepest part of Jesse Ewing Canyon. At mile 58.9, cross the drainage divide and enter the Clay Basin gas field after crossing

the Uinta-Sparks fault zone at the northern structural edge of the Uinta uplift. Stay left at the intersections and cross Red Creek as the road changes to pavement at mile 64.1. The road ahead follows the Uinta fault, separating, at left, Precambrian metamorphic rocks of the Red Creek Quartzite, and to the right, tilted Cretaceous (Mesa Verde Group) to Eocene (Wasatch Formation on skyline) sedimentary rocks of the Green River Basin. Here is the greatest throw of any structure in the Uintas, and it presumably once hosted the highest part of the Laramide range to the south, before collapse. Cross back onto the gravel road at the Wyoming state line at mile 72.1, and at mile 75.0, turn left on Hwy 191, crossing the Utah state line through a hogback of the Mesa Verde Group. Continue through the hogback of Jurassic Glen Canyon (Nugget) Sandstone to Dutch John at mile 83.6. Turn left and proceed to the Little Hole overlook turnout on the right side of the road at mile 89.0. Stop 2.4 (Mile 89.0) This stop at Little Hole, a small park-like area within lower Red Canyon 10 km downstream of Flaming Gorge Dam, focuses on the anomalous Qag5 fill and features a short optional hike to an outcrop containing Lava Creek B tephra. Highly variable surficial deposits are preserved at Little Hole. The four lowest deposits are 5–10 m thick, typical Green River gravels that are each associated with a distinct terrace (Fig. 13). In contrast, the sand-dominated Qag5 deposit at Little Hole and just downstream at Devils Hole are at least 60 m thick

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(the base of the deposit is obscured here), and the regular Green River terrace gravels are inset into this older fill. Qag5 here is generally a medium-scale, cross-bedded, pebbly sand with clastsupported, relatively immature, pebble-boulder gravel lenses of locally derived UMG clasts (Fig. 13). The middle-upper exposed deposit generally coarsens upward, and pebbles of brown and tan quartzite, Paleozoic, and Mesozoic bedrock clasts typical of mainstem Green River deposits appear as a terrace-gravel cap near the top. Laterally equivalent, coarse, local piedmont gravels of UMG cobbles and boulders grade to the higher preserved terrace level near the top of the deposit. At both Little Hole and Devils Hole, thin- to medium-scale lenticular beds of fine grained, in most cases reworked, tephra lie near the preserved top of the deposit (Fig. 13). Geochemical analysis of these beds indicates the presence of Lava Creek B ash (Mike Perkins, 2004, personal commun.). Thus, the upper strata of this anomalous sandy deposit were emplaced soon after the Lava Creek B eruption at ca. 640 ka in the middle Pleistocene (Lanphere et al., 2002). Hansen (1965) recognized that these deposits were in some ways more similar to the Browns Park Formation than to typical Pleistocene Green River deposits, and he understandably mapped them as the heterogeneous Browns Park basin fill. We have reinterpreted this middle Pleistocene deposit as sediment of the Green River and local tributaries deposited in the accommodation space upstream of the hypothesized landslide dam in lower Red Canyon. The Qag5 terrace here is at approximately the same elevation as the reworked top of the landslide toe. Results of sand petrography of the Qag5 deposit at Little Hole support this interpretation, with QFL point-count data of samples indicating a composition that matches Pleistocene Green River deposits rather than examples of basin fill in western Browns Park (Counts, 2005). Directions to Red Canyon Lodge Return to Dutch John (mile 94.4), turn left, and head south on Hwy 191, cross Flaming Gorge Dam, climb out of canyon, and turn right at the intersection with Rt. 44 at mile 102.5. Continue west to mile 106.3 then turn right on Red Canyon Road to Red Canyon Lodge. DAY 3—RED CANYON LODGE TO SALT LAKE CITY—THE NORTHERN UINTA MOUNTAINS Introduction The final day of this trip (driving distance ≈225 mi [360 km]) focuses on the latest Pleistocene glacial geology and geomorphology of the north slope of the Uintas, returning to Salt Lake City on I-80 through Echo and Parley’s canyons. Directions to Stop 3.1 From Red Canyon Lodge (mile 0.0), backtrack to Rt. 44, and turn right. At 9.4 mi, the road descends into the deeply incised Carter Creek drainage, which carried meltwater away from the northeasternmost glaciers in the Uintas. A 12.1 mi, turn left on

Figure 13. (A) Composite surveyed cross section from Little Hole (modified from Counts, 2005). The Qag5 deposit is anomalously sandy and thick (note that vertical exaggeration is the same as in Figure 11B). (B) Roadcut of Qag5 deposit at Little Hole is primarily fluvial sand containing lenses of UMG-dominated cobble gravel derived from local catchments. Typical mainstem and piedmont terrace gravels overlie and are inset into the sandy fill, and the Lava Creek B tephra lies near the top. Qag—Quaternary alluvial gravel of mainstem drainages; Qagp—Quaternary alluvial gravel of piedmont drainages.

the Sheep Creek Geologic Loop (FR 218). Views of Leidy Peak and the cirques at the head of Carter Creek appear at mile 15.2. Turn left at mile 15.3 on FR 221, and immediately bear right, following the signs for the Ute Tower. The fire tower is accessed via a short (1.3 mi) side road (FR 005) on the left at mile 16.6. Stop 3.1—Ute Tower (Mile 16.6) The Ute Tower provides a tremendous overview of the eastern High Uintas that is an illustrative way to start the third day of this field trip. Due south of the tower is the easternmost alpine summit of the High Uintas, Leidy Peak, named for nineteenth century paleontologist Joseph Leidy (Figs. 14 and 15). Leidy Peak rises above the extensive, gently sloping, summit flats known as “bollies” at this end of the range. In his monograph on the geomorphology of the northern Uinta, Bradley (1936) interpreted this summit surface as a remnant of a pediment formed prior to mountain uplift during a long period of erosion under a semiarid climate. However, recent research and modeling on

New Quaternary research in the Uinta Mountains

71 Figure 14. Map showing the locations of stops on the eastern north slope of the Uintas overlain on a 30-m digital elevation model. The dark line marks the field trip route (east to west); the stops are labeled as in Figure 1. Leidy Peak (LP), Deep Creek (DC), and The Narrows (see text) are highlighted. Box A marks the location of Figure 17; Box B marks Figure 20. Outlines of glaciers at the Smiths Fork maximum (Munroe, 2001; Laabs, 2004) are shown with the cross-hatched pattern. Black highlights the high elevation summit flats of the eastern Uintas (from Munroe, 2005). Glacier names: 1—North Fork Ashley Creek; 2—East Fork Carter Creek; 3— West Fork Carter Creek; 4—South Fork Sheep Creek; 5—Middle Fork Sheep Creek; 6—Burnt Fork; 7—Thompson Creek; 8—Middle Fork Beaver Creek; 9—West Fork Beaver Creek; 10—Henrys Fork; 11—East Fork Smiths Fork; 12—Uinta; 13—Whiterocks; 14—Dry Fork; 15—South Fork Ashley Creek.

summit flats in the Wind River Range, ~200 km to the north (Anderson, 2002; Small et al., 1997), has suggested that summit flats can form in place through periglacial hillslope processes and can result solely from the difference in rate between periglacial and glacial erosion. Given this theory, Munroe (2005) investigated the distribution of summit flats throughout the Uintas and found that these surfaces are considerably more extensive at the eastern end of the range (Fig. 14). He speculated that the paucity of summit flats at the western, upwind, end of the Uintas may reflect more effective glacial erosion there over the course of the Quaternary. This idea is corroborated by reconstructions of Smiths Fork glaciers by Oviatt (1994), Munroe (2001), and Laabs (2004), which indicate that the western Uintas were more extensively glaciated during at least the last few glacial cycles. The contrasting morphologies of glaciated and nonglaciated valleys on the north side of Leidy Peak provide a good example of a “glaciation threshold” conditioned by local snow redistribution. Deep Creek (DC on Fig. 14) begins on the northeastern shoulder of Leidy Peak at an elevation of ~3120 m and flows eastward for 1.5 km before turning abruptly northward and descending almost 250 m in 3 km. Throughout this length, the Deep Creek valley exhibits the typical V-shaped profile of a fluvial system, and nowhere do deposits typical of glaciation interrupt the course of the river. In contrast, the East Fork Carter Creek, located immediately west of Deep Creek, originates in the easternmost cirque on the north slope. Red Lake (3015 m) in this cirque is overdeeped by at least 17 m (Pettengill, 1996) and is flanked along its western and southern sides by a steep headwall that rises 200 m to meet the bollies on the north side of Leidy Peak. The headwaters of both drainages lie at similar elevations and are only 4.5 km apart,

yet Deep Creek was apparently never glaciated. This relationship underscores the sensitivity of small glaciers to local variations in snow accumulation. The East Fork Carter Creek lies immediately downwind of the summit flats flanking Leidy Peak (Fig. 15). Snow blown from this surface—the easternmost extensive alpine upland in the Uintas—nourished the accumulation area of the East Fork Carter Creek glacier, while Deep Creek, arising at a similar elevation, but slightly farther east of Leidy Peak, did not receive as large a volume of drifting snow.

Figure 15. View southward from Ute Tower. Leidy Peak is the easternmost alpine summit in the Uinta Mountains; Red Lake cirque was the accumulation area for the easternmost glacier in the northern Uintas. The Smiths Fork (SF) and Blacks Fork (BF) age moraines are visible as prominent benches, highlighted by the evening light.

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End moraines of the Smiths Fork and Blacks Fork glaciations are also visible to the south and southwest from Ute Tower, where Smiths Fork–age glaciers in the East and Middle Forks of Carter Creek coalesced to form a piedmont lobe (Fig. 15). The prominent terminal moraine runs ~E-W for several kilometers at an elevation of 2775 m, enclosing Tepee Lakes as well as other unnamed ponds (Leidy Peak quadrangle). The moraine is typically sharp-crested, with a steep frontal slope that rises up to 60 m above the older till deposits to the north. The analogous terminal moraine in the West Fork Carter Creek of Smiths Fork–age is over 35 m high and fronts a chaotic zone of hummocky topography including numerous water-filled closed depressions. Older glacial deposits are also well preserved in the Carter Creek region. Immediately in front of the Smiths Fork terminal moraine is a lower relief, Blacks Fork–age, drift sheet that has a slightly hummocky surface and contains only isolated kettles (Fig. 15). Blacks Fork till is also present beyond the Smiths Fork–age moraine along the West Fork Carter Creek. This surface also displays subdued hummocks, partially filled kettles, and poorly drained areas. No evidence for pre–Blacks Fork till is found in the Carter Creek area. Instead, outcrops of bedrock ledge and flaggy surface clasts indicate that surficial deposits beyond the Blacks Fork moraines were derived from in situ weathering of bedrock. Directions to Stop 3.2 Descend the road from Ute Tower and turn left on FR 221 (mile 19.2). End moraines of the West Fork Carter Creek are visible straight ahead at mile 22.0, after leaving Half Moon Park. Continue past the turn for Browne Lake at mile 22.7, passing several bedrock-cored hills similar to the one on which Ute Tower is located. At mile 28.8, turn left on FR 001 following the signs for Spirit Lake. The road passes around the outside of Hickerson Park before rising toward the terminal moraine complex in the South Fork Sheep Creek valley. At mile 29.4, the Smiths Fork–age terminal moraine forms the skyline ridge straight ahead, while the unglaciated North Fork Ashley Creek is visible on the right. The road climbs onto subdued Blacks Fork–age till at mile 30.2 and reaches the front of the Smiths Fork–age terminal moraine at mile 30.7. Park in the turnout on the left before the road curves to the right. Stop 3.2—Smiths Fork–Age End Moraine, Middle Fork Sheep Creek (Mile 30.7) During the Smiths Fork glaciation, glaciers in the Middle and South forks of Sheep Creek were larger than those in the Carter Creek drainages (~25 km2 versus <20 km2). However, glaciers in all of these valleys, more than those in other drainages of the Uintas, tended to broaden out in their terminal zones, forming piedmont lobes. The uniqueness of these glaciers is apparently due to the structural geology in this part of the eastern north slope. Because the glacial valleys dip less steeply than the Uinta Mountain Group in this area, the valleys become progressively shallower to the north, allowing the glaciers to broaden into piedmont lobes

on the steeply dipping north flank of the Uinta arch. Furthermore, all four of these glaciers terminated upslope from the Paleozoic hogback, so they did not encounter this prominent topographic obstacle as glaciers farther west did. Divergent ice flow across the bedrock dipslope assisted ice stagnation within the terminal zones by increasing the rate of ice thinning toward the margin. As a result, Smiths Fork–age terminal moraines along the forks of Sheep Creek and Carter Creek incorporated large amounts of buried ice and consequently formed areas of hummocky topography. Deposits of at least two glaciations are present along the forks of Sheep Creek, straddling the boundary between the Whiterocks Lake and Phil Pico Mountain 7.5′ quadrangles (Fig. 14). Immediately south of Hickerson Park is an unglaciated area of bedrock-controlled topography featuring linear strike ridges of quartzite, colluvial basins controlled by bedrock sills, and a conspicuous absence of subrounded quartzite erratics. Moving southward, this thin, locally derived regolith gives way to a heavily eroded till unit, identifiable by its abundance of subrounded erratic boulders. No moraine marks the outer limit of this drift sheet, but the appearance of the first erratics south of the Lodgepole Creek crossing along FR 001 is conspicuous, and the position and elevation (~2810 m) of this deposit suggest that it is analogous to the Blacks Fork–age deposits along Carter Creek. The main focus of this stop is the Smiths Fork–age terminal moraine, found at an elevation of 2875 m along FR 001. Accessible, extensive exposures of Smiths Fork–age till are rare in the Uintas, and the prominent roadcut on the west side of the road is exceptional because of its scale; its freshness is ensured by road maintenance. The exposure reveals the diamicton typically found in Smiths Fork–age terminal moraines in the northern Uintas, which were apparently produced by a wide variety of sedimentary mechanisms. Large surface boulders (>50 cm in diameter), some partially buried by eolian silt, were likely deposited directly from a supraglacial position at the ice margin, either due to wasting of the underlying ice or by direct rolling and sliding to the glacier terminus. The sandy till present below the silt cap is unstratified and is interpreted as meltout till deposited at the ice margins. Fabric is difficult to evaluate in most Smiths Fork–age till due to the paucity of elongate clasts. Basal till is rarely exposed at the surface in the northern Uintas, although diamicton containing striated bullet clasts exposed by landsliding along the Little West Fork Blacks Fork may represent basal till. Overconsolidated till with well developed columnar jointing is exposed along the West Fork Beaver Creek upstream from the Smiths Fork–age terminal moraine complex. A unique feature of the exposure on Spirit Lake Road is the slabs of bedrock exposed in the sandy till matrix. Is this evidence for wholesale entrainment of bedrock ledges from points further south in the basin? Or is this moraine bedrock-cored? We will spend some time trying to answer this question before moving on to our next stop. Directions to Stop 3.3 Retrace the route back to FR 221, and turn left at the intersection (mile 32.6). At mile 35.3, continue straight on FR

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Figure 16. Hole in the Rock and the Beaver Creek area viewed from Stop 3.3 near Lonetree, Wyoming.

221 and descend Birch Canyon to the north. Red Pine Shale is exposed along the road near mile 35.7, and the valley narrows as the Madison Limestone is encountered near mile 37.4. Views of the Smiths Fork–age outwash surface graded to the terminal moraine in the Burnt Fork valley open up on the left at mile 42.4. Turn left on Hwy 414 toward Lonetree, Wyoming, at mile 45.9. At mile 46.6, the road curves left and crosses the Burnt Fork outwash surface, and at mile 47.8, passing the marker for the 1825 Rendezvous, the forested hills of the Bald Range are visible in the middle distance. Enter Uinta County at mile 50.8, with views of the Henrys Fork on the right at mile 52.1. Turn left at mile 56.4 on CR 290 toward the Henrys Fork, and then turn left on CR 291 at mile 57.4. The rounded sage-covered upland on the right after the turn is underlain by pre–Black Fork-age till. Continue straight through the intersection at mile 58.8, and park at the entrance to the gravel pit on the right at mile 59.9. Stop 3.3—Gravel Pit South of Lonetree, Wyoming (Mile 59.9) This stop provides a spectacular view of the north face of the High Uintas and allows us to review the alluvial stratigraphy below the glacial termini of the Henrys Fork drainage. Lonetree is situated at the northern end of an expansive compound outwash surface formed by the combined meltwater from the Henrys Fork, West Fork Beaver Creek, and Middle Fork Beaver Creek glaciers (Fig. 14). To the south, the prominent strike-ridge of the Mississippian Madison Limestone is visible (Fig. 16), forming “Hole in the Rock” at the left. The central gap in the limestone was formed by meltwater from the Middle Fork Beaver Creek; Smiths Fork–age moraines in this valley are hidden behind the limestone ridge (Fig. 17). The most westerly (and widest) gap in the limestone is partially blocked by end moraines of the West Fork Beaver Creek (Fig. 17). Bryant (1992) mapped deposits of the Smiths Fork, Blacks Fork, and pre–Blacks Fork glaciations in this moraine complex. More recently, Douglass (2000) examined the morphology of these moraines and their weathering characteristics. His results support the mapping of Bryant, and further suggest that the Smiths Fork–age moraine is a compound feature, possibly representing ice advances during MIS 2 and 4. To the east-southeast of Lonetree lies the Bald Range, a complex upland landscape of random hills and local internal drainage (Fig. 14). Bradley (1936) interpreted these sediments as deposits of pre–Blacks Fork glaciers that advanced from the northern Uintas presumably before the modern Beaver

Creek–Henrys Fork valley was developed. However, poorly consolidated sediment of the Tertiary Wasatch Formation is exposed along the west side of the Bald Range, and diagnostic evidence of a glacial origin of the capping sediments is lacking. Gibbons and Hansen (1980) argued for a reinterpretation of the Bald Range as landslide deposits, and Gibbons (1986) mapped the entire Bald Range as landslide deposits on the “Surficial materials map of the Evanston 30′ × 60′ quadrangle.” The exact manner in which mass wasting produced the modern form of the Bald Range is unclear. However, fluvial erosion on three sides (Beaver Creek–Henrys Fork on the west, Henrys Fork on the north, and Burnt Fork on the west) clearly isolated this pile of weak sediments, which may have shifted and failed in response to a seismic trigger. The gravel pit at this site also provides an opportunity to review the alluvial stratigraphy below the glacial termini of the Henrys Fork drainage, which has been mapped at 1:24,000 scale in the development of a nonglacial Quaternary stratigraphic

Figure 17. Moraines of the Middle Fork and West Fork Beaver Creek and their compound outwash surface south of Lonetree, Wyoming (see Fig. 14 for location). Prominent east-west ridge is a hogback of the Mississippian Madison Limestone. The Middle Fork Beaver Creek glacier terminated upslope from the limestone during the Smiths Fork Glaciation, while the West Fork glacier advanced slightly beyond the hogback. Moraines corresponding to the Blacks Fork Glaciation, and till of pre–Blacks Fork age, are present farther north. The compound outwash surface extends over 10 km northward to merge with outwash deposits in the Henrys Fork river valley.

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framework for the northeastern Uinta Mountains (Counts, 2005). The Henrys Fork valley contains nine distinct mainstem gravels, six piedmont gravels, and landslide deposits (Fig. 18). Gravels on the Henrys Fork are grouped as older, higher remnant gravels that cannot be directly linked to glacial units and as younger gravels that can be traced from glacial moraines, through outwash plains, to stream-valley gravels with terraces formed upon them. This example of Qag3 (Blacks Fork equivalent) gravels are representative of the clast-supported, cobble gravel of the Henrys Fork, derived mostly from the UMG and Paleozoic limestone units. Near moraines, gravels are thicker but they quickly thin downstream and lie on planar bedrock straths, and so form strath terraces that converge downstream. Besides Qag1, which is correlated to the dated Smiths Fork glaciation, Henrys Fork terraces Qag2 and Qag3 are tentatively correlated to relatively well dated Wind River terraces (Counts, 2005). Incision rate estimates for the Henrys Fork from this correlation are 80–110 m/my over the late Pleistocene (Table 1). Extrapolating a linear incision rate (maximum long-term estimate) suggests that the oldest gravels on the Henrys Fork were deposited in the early Pleistocene. Directions to Stop 3.4 Retrace the route back to the last intersection (mile 61.1) and turn left on CR 263. Bear right at mile 61.5, following signs for the Henrys Fork. Cross back into Utah at mile 67.2, and pass the Forest Boundary at mile 67.9, with the Smiths Fork–age terminal moraine visible straight ahead. The road enters the moraine near mile 68.6 and passes through an area of high-relief hummocky topography before descending the proximal slope of the moraine at mile 69.3. Cross the Henrys Fork at mile 69.7, with a view of Gilbert Peak upstream, and turn right into the Quarter Corner Trailhead at mile 71.3 for lunch (bathroom, picnic tables). After lunch, return to the main road and continue to the right, taking a hard right at the intersection with FR 017 at mile 71.8. The road climbs up and over the west side of the terminal moraine and crosses Dahlgreen Meadow (mile 74.0) before climbing a series of switchbacks up the south face of Red Mountain. Turn right on

Figure 18. Composite cross-valley profile showing landscape position of major Henrys Fork terrace gravels (modified from Counts, 2005). Ts—Tertiary basin-fill sediment. Qag4 gravel was not identified along the Henrys Fork, but it exists along the tributary Beaver Creek and Burnt Fork. Qag—Quaternary alluvial gravel of mainstem drainages; Qagp—Quaternary alluvial gravel of piedmont drainages; Qal—alluvium of active channels and floodplains.

FR 155 at mile 75.1 and park at the entrance to FR 132 on the right (mile 75.3). From the vans, we will hike for ~15 min up FR 132 to the summit of Red Mountain. Stop 3.4—Red Mountain (Mile 75.3) Red Mountain provides a spectacular view of the central High Uintas (Fig. 19). Several of the most prominent peaks visible from this point are named for famous geologists, including Kings Peak (4136 m—due south), Gilbert Peak (4110 m—SSE), and Mount Powell (4023 m—SSW). The broad alpine basin west of Gilbert Peak formed the accumulation area for a glacier in the Henrys Fork that covered 76 km2 at the LGM. The glacier, which was over 24 km long, advanced beyond the hogback of Madison Limestone at The Narrows (visible 5 km upvalley) and formed a hummocky moraine east of Red Mountain (Figs. 14, 19, and 20). Because the cross-sectional area of the valley within The Narrows is 25% less than that immediately upstream, and because the Henrys Fork glacier overtopped The Narrows by only 20–50 m, relatively modest downwasting during deglaciation would have separated the ice in the terminal zone from the active ice in the valley to the south. This geometry likely caused the terminal zone of the Henrys Fork

Figure 19. View southward from Red Mountain. Several major peaks of the High Uintas are visible in the distance. The Narrows is a prominent valley constriction formed where the Henrys Fork passes through the Madison Limestone hogback. The Smiths Fork–age terminal moraine of the Henrys Fork glacier is visible as a forested ridge in the foreground. Dahlgreen Meadow marks the site of a former ice-marginal lake, formed when the Henrys Fork glacier impinged upon the base of Red Mountain, damming an ice-marginal drainage.

New Quaternary research in the Uinta Mountains glacier to stagnate at the start of deglaciation, even when glaciers in neighboring valleys were able to keep pace with ameliorating conditions by actively retreating upvalley. Interestingly, Atwood (1909) mapped this terminal moraine as a Blacks Fork equivalent and located the Smiths Fork–age terminal moraine near Alligator Lake, 6 km upvalley from The Narrows. However, the overall fresh form of this massive moraine loop, its steep frontal slope, and its dramatically kettled surface argue that it was formed during the Smiths Fork glaciation. Furthermore, outwash graded to this moraine merges with Smiths Fork–age outwash from the West and Middle forks of Beaver Creek. Thus, the moraine near Alligator Lake is most likely a rock-cored recessional feature formed when the retreating Smiths Fork–age glacier was temporarily pinned on a bedrock high. The advance and retreat of the Henrys Fork glacier also impacted the fluvial system through formation of ice-marginal channels and ice-dammed lakes. Ice-marginal channels containing underfit streams are common components of the north slope landscape. Dahlgreen Creek, which runs alongside the western lateral moraine of the Henrys Fork below Red Mountain, is a good example (Fig. 19). In some cases, lateral moraines redirected ice-marginal drainages into other watersheds, for instance upstream from The Narrows on the east side of the Henrys Fork where the voluminous right lateral moraine contained a stagnant ice body that drained eastward through Bullocks Park into the Beaver Creek system along the route followed by FR 058. Meltwater derived from this stagnant ice mass appears to have helped carve a short canyon along Fallon Creek (Hole in the Rock quadrangle). Ice-marginal channels were also vulnerable to impoundment during periods of ice advance, leading to the formation of temporary ice-marginal lakes. The willow-filled meadow just south of Red Mountain was likely the bed of an ice-marginal lake during times when the Henrys Fork glacier blocked the northward drainage of Dahlgreen Creek (Fig. 19). Finally, the summit of Red Mountain was mapped by Bryant (1992) as the Tertiary Bishop Conglomerate (see Stop 2.1). Red Mountain and the surface extending southward to increasingly higher elevations along the west side of the Henrys Fork are remnants of Bradley’s (1936) Gilbert Peak erosion surface mantled by Bishop Conglomerate. Although the exact mechanism by which this extensive gently sloping landform was generated is unclear, Red Mountain appears to have been isolated by slumping along its southern and eastern faces, perhaps in response to slope steepening due to glacial erosion by the Henrys Fork glacier. Directions to Stop 3.5 After walking back to the vans, turn around and head south to the intersection with FR 017 (mile 75.6). Turn right to continue west into the East Fork Smiths Fork drainage. At the next intersection (mile 79.4), turn right toward Mountain View, Wyoming. The slope directly across the road from the stop sign is the distal side of the right lateral moraine in the East Fork Smiths Fork valley. Cross the Wasatch-Cache National Forest boundary at mile 83.1, and at mile 84.3, turn left on CR 285. Cross back into the

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Figure 20. Location of the Smiths Fork–age moraine (dashed) in the Henrys Fork (see Fig. 14). The Smiths Fork–age glacier in the Henrys Fork advanced northward beyond the prominent valley constriction of The Narrows, formed by the Mississippian Madison limestone. The piedmont lobe formed by the glacier stagnated during deglaciation when the glacier surface lowered, dramatically reducing the cross-sectional area feeding the terminal zone. A smaller area of stagnant ice developed on the southeastern side of The Narrows; this zone drained eastward, depositing outwash on the south side of the limestone. Dahlgreen Creek (an ice-marginal drainage) was temporarily impounded by the glacier terminus, forming a lake in the basin south of Red Mountain (black dot). The route of the field trip southwest from Lonetree, Wyoming, on County Road 263, and westward on Forest Road 058, is also shown (black line).

forest at mile 86.0, with the terminal moraine of the East Fork Smiths Fork visible to the left of center. Outwash graded to this moraine is visible along the left side of the road at mile 86.8. Cross the river and park on the right at mile 86.9. Stop 3.5—Smiths Fork Type Locality (Mile 86.9) The terminal moraine considered by Bradley to be the Smiths Fork glaciation type locality is present just south of the

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road at the bridge, on either side of the stream. The western half of the terminal moraine is vegetated by lodgepole pine and stands in stark contrast to the sagebrush-covered outwash surface. The eastern half of the moraine (at least the part visible from this vantage) forms an unforested slope rising steeply above the river. The crests of both moraines are <1 km wide and marked by slightly sinuous subparallel ridges standing <10 m above their surroundings. Large erratic boulders of quartzite are rare, but several were located and sampled for cosmogenic surface exposure dating during the 2003 and 2004 field seasons. Results of these analyses, which should help constrain the timing of the local LGM in the northern Uintas, are pending. Directions to Salt Lake City Turn around and head back toward Mountain View, Wyoming, turning left at the stop sign (mile 89.5). Continue straight at the intersection with the paved road (mile 96.7) and turn left at the junction with Hwy 414 in Mountain View (mile 103). Pass through Mountain View and continue straight at the intersection in Urie (mile 106.8). Cross the Blacks Fork at mile 109.1 and enter I-80 westbound at mile 109.7. Follow I-80 west to Evanston (mile 144) and on to Salt Lake City (~mile 224). ACKNOWLEDGMENTS We thank the employees of the Ashley and Wasatch-Cache National Forests, particularly D. Koerner and S. Ryberg, for logistical support through numerous field seasons. Funding for our research has been provided by grants from sources including the National Science Foundation, the Geological Society of America, Sigma Xi, the Desert Research Institute, the American Alpine Club, the University of Wisconsin–Madison, Middlebury College, and Utah State University. Finally, we thank the many students and technicians who provided assistance in the field and lab. REFERENCES CITED Anderson, R.S., 2002, Modeling the tor-dotted crests, bedrock edges, and parabolic profiles of high alpine surfaces of the Wind River Range, Wyoming: Geomorphology, v. 46, p. 35–58, doi: 10.1016/S0169-555X(02)00053-3. Andrews, E.D., 1986, Downstream effects of Flaming Gorge Reservoir on the Green River, Colorado and Utah: Geological Society of America Bulletin, v. 97, p. 1012–1023. Atwood, W.W., 1909, Glaciation of the Uinta and Wasatch Mountains: U.S. Geological Survey Professional Paper 61, 96 p. Barnhardt, M.L., 1973, Late Quaternary geomorphology of the Bald Mountain area, Uinta Mountains, Utah [M.S. thesis]: Salt Lake City, University of Utah, 109 p. Bauer, M.S., 1985, Heat flow at the Upper Stillwater Dam site, Uinta Mountains, Utah [M.S. thesis]: Salt Lake City, University of Utah, 94 p. Bradley, W.A., 1936, Geomorphology of the north flank of the Uinta Mountains: U.S. Geological Survey Professional Paper 185-I, p. 163–199. Bradley, M.D., 1995, Timing of the Laramide rise of the Uinta Mountains, Utah and Colorado, in Jones, R.W., ed., Resources of southwestern Wyoming: Wyoming Geological Association 1995 Field Conference Guidebook, p. 31–44. Brink, M.S., and Schmidt, J.C., 1996, The Duchesne River channel: A geomorphic history, 1875 to 1995: Logan, Utah State University, Final Report to the Utah Division of Wildlife Resources, 47 p.

Bryant, B., 1992, Geologic and structure maps of the Salt Lake City 1° × 2° quadrangle, Utah and Wyoming: U.S. Geological Survey Miscellaneous Geological Investigations Map I-1997, scale 1:125,000. Butler, B.S., Laughlin, G.F., and Heikes, V.S., 1920, The ore deposits of Utah: U.S. Geological Survey Professional Paper 111, 672 p. Chadwick, O.A., Hall, R.D., and Phillips, F.M., 1997, Chronology of Pleistocene glacial advances in the central Rocky Mountains: Geological Society of America Bulletin, v. 109, no. 11, p. 1443–1452, doi: 10.1130/00167606(1997)109<1443:COPGAI>2.3.CO;2. Counts, R.C., 2005, The Quaternary stratigraphy of the Henrys Fork and western Browns Park, northeastern Uinta Mountains, Utah and Wyoming [M.S. Thesis]: Logan, Utah State University, 165 p. Condie, K.C., Lee, D., and Farmer, G.L., 2001, Tectonic setting and provenance of the Neoproterozoic Uinta Mountain and Big Cottonwood groups, northern Utah: Constraints from geochemistry, Nd isotopes, and detrital modes: Sedimentary Geology, v. 141-142, p. 443–464, doi: 10.1016/ S0037-0738(01)00086-0. Dehler, C.M., Porter, S., De Grey, L.D., and Sprinkel, D.A., 2006, The Neoproterozoic Uinta Mountain Group revisited: A synthesis of recent work on the Red Pine shale and related undivided clastic strata, Northeastern Utah, in Link, P.K., and Lewis, R., eds., Proterozoic basins of the northwestern U.S.: Society for Sedimentary Geology (SEPM) Special Volume (in press). Douglass, D.C., 2000, Glacial history of the West Fork of Beaver Creek, Uinta Mountains, Utah [M.S. thesis]: Madison, University of Wisconsin, 64 p. Epis, R.C., and Chapin, C.E., 1975, Geomorphic and tectonic implications of the post-Laramide, late Eocene erosion surface in the southern Rocky Mountains, in Epis, R.C., and Chapin, C.E., eds., Cenozoic history of the southern Rocky Mountains: Geological Society of America Memoir 144, p. 45–74. Gaeuman, D.A., Schmidt, J.C., and Wilcock, P.R., 2005, Complex channel responses to changes in stream flow and sediment supply on the lower Duchesne River, Utah: Geomorphology, v. 64, p. 185–206, doi: 10.1016/ j.geomorph.2004.06.007. Gibbons, A.B., and Hansen, W.R., 1980, Origin of the Bald Range, southeast [sic] Wyoming: U.S. Geological Survey Professional Paper Report P1175, 273 p. Gibbons, A.B., 1986, Surficial materials map of the Evanston 30′ × 60′ quadrangle, Uinta and Sweetwater Counties, Wyoming: U.S. Geological Survey Coal Investigations Map C-103, scale 1:100,000. Gilmer, D.R., 1986, General geology, landsliding, and slope development of a portion of the north flank of the Uinta Mountains, south-central Uinta County, Wyoming [M.S. thesis]: Laramie, University of Wyoming, 93 p. Grams, P.E., and Schmidt, J.C., 1999, Geomorphology of the Green River in the eastern Uinta Mountains, Dinosaur National Monument, Colorado and Utah, in Miller, A.J., and Gupta, A. eds., Varieties of fluvial forms: John Wiley and Sons, p. 81–111. Grogger, P.K., 1974, Glaciation of the High Uintas Primitive Area, Utah, with emphasis on the northern slope [Ph.D. thesis]: Salt Lake City, University of Utah, 209 p. Hansen, W.R., 1965, Geology of the Flaming Gorge area, Utah, Colorado, and Wyoming: U.S. Geological Survey Professional Paper P-490, 196 p. Hansen, W.R., 1969, Development of the Green River drainage system across the Uinta Mountains, in Lindsay, J.O., ed., Geologic guidebook of the Uinta Mountains, Utah’s maverick range: Intermountain. Association Geologists and Utah Geological Society, 16th Annual Field Conference 1969, p. 93–100. Hansen, W.R., 1984, Post-Laramide tectonic history of the eastern Uinta Mountains, Utah, Colorado, and Wyoming: The Mountain Geologist, v. 21, p. 5–29. Hansen, W.R., 1986, Neogene tectonics and Geomorphology of the Eastern Uinta Mountains in Utah, Colorado, and Wyoming: U.S. Geological Survey Professional Paper 1356, 78 p. Hansen, W.R., Carrara, P.E., and Rowley, P.R., 1982, Geologic map of the Canyon of Lodore North Quadrangle, Moffat County, Colorado: U.S. Geological Survey map GQ-1568, scale 1:24,000. Hayden, F.V., 1871, Preliminary Report of the United States Geological Survey of Wyoming and portions of contiguous territories: Washington, U.S. Government Printing Office, 511 p. King, C., 1878, Systematic Geology, United States Geographical Exploration of the Fortieth Parallel: Washington, U.S. Army Engineering Department, 803 p.

New Quaternary research in the Uinta Mountains Kowallis, B.J., Christiansen, E.H., Balls, E., Heizler, M.T., and Sprinkel, D.A., 2005, The Bishop Conglmerate ash beds, south flank of the Uinta Mountains, Utah: Are they pyroclastic fall beds from the Oligocene ignimbrites of western Utah and eastern Nevada?, in Dehler et al., eds., Uinta Mountain geology: Salt Lake City, Utah, Utah Geological Association Publication 33 (in press). Laabs, B.J.C., 2004, Late Quaternary glacial and paleoclimate history of the southern Uinta Mountains, Utah. [Ph.D. thesis]: Madison, University of Wisconsin, 162 p. Laabs, B.J.C., and Carson, E.S., 2005,The glacial geology of the Uinta Mountains, in Dehler et al., eds., Uinta Mountain geology: Salt Lake City, Utah, Utah Geological Association Publication 33 (in press). Lanphere, M.A., Champion, D.E., Christiansen, R.L., Izett, G.A., and Obradovich, J.D., 2002, Revised ages for tuffs of the Yellowstone Plateau volcanic field, assignment of the Huckleberry Ridge Tuff to a new geomagnetic polarity event: Geological Society of America Bulletin, v. 114, no. 5, p. 559–568, doi: 10.1130/0016-7606(2002)114<0559: RAFTOT>2.0.CO;2. Larsen, I.J., Schmidt, J.C., and Martin, J.A., 2004, Debris fan reworking during low magnitude floods in the Green River canyons of the eastern Uinta Mountains, Colorado and Utah: Geology, v. 32, p. 309–312, doi: 10.1130/G20235.1. Lenfest, L.W., Jr., and Ringen, B.H., 1985, Streamflow and suspended-sediment discharge from two small watersheds in southwestern Wyoming and northeastern Utah, 1984: U.S. Geological Survey Open File Report 85-161, 35 p. Martin, J.A., 2000, Debris-flow activity in Canyon of Lodore, Colorado: Implications for debris-fan formation and evolution [M.S. Thesis]: Logan, Utah State University, 155 p. Mears, B., Jr., 1993, Geomorphic history of Wyoming and high-level erosion surfaces, in Snoke, A.W., Steidtmann, J.R., and Roberts, S.M., eds., Geology of Wyoming: Geological Survey of Wyoming Memoir 5, p. 608–627. Munroe, J.S., 2001, Late Quaternary history of the northern Uinta Mountains, northeastern Utah [Ph.D. thesis]: Madison, University of Wisconsin, 398 p. Munroe, J.S., 2005, Investigating the spatial distribution of summit flats in the Uinta Mountains of northeastern Utah, U.S.A., in Munroe, J.S., Laabs, B.J.C., and Carson, E.C., eds., Quaternary landscape change and modern process in western North America: Geomorphology, special issue (in press). Nelson, A.R., and Osborn, G.D., 1991, Quaternary history of some southern and central Rocky Mountain basins: Northwestern Uinta Basin, in Morrison, R.B., ed., Quaternary nonglacial geology: Conterminous U.S.: Boulder, Colorado, Geological Society of America, Geology of North America, v. K-2, p. 432–440. Osborn, G.D., 1973, Quaternary geology and geomorphology of the Uinta Basin and the south flank of the Uinta Mountains, Utah [Ph.D. thesis]: Berkeley, University of California, 266 p. Oviatt, C.G., 1994, Quaternary geologic map of the upper Weber River drainage basin, Summit County, Utah: Utah Geological and Mineral Survey Report 156, scale 1:50,000. Palmquist, R., 1983, Terrace chronologies in the Bighorn Basin, Wyoming, in Boberg, W.W., ed., Geology of the Bighorn Basin, Wyoming Geological Association Guidebook: Casper, Wyoming Geological Association, p. 217–231. Pederson, J.L., Mackley, R.D., and Eddleman, J.L., 2002, Colorado Plateau uplift and erosion—amounts and causes evaluated with GIS: GSA Today, v. 12, no. 8, p. 4–10, doi: 10.1130/1052-5173(2002)012<0004: CPUAEE>2.0.CO;2. Pederson, J.L., 2004, Drainage integration as a first-order control on the erosional exhumation of the Interior West—The example of the Green River and the Uinta Mountains: Geological Society of America Abstracts with Programs, v. 36, no. 5, p. 118. Pederson, J.L., and Hadder, K.W., 2005, Revisiting the classic conundrum of the Green River’s integration through the Uinta uplift, in Dehler et al., eds., Uinta Mountain geology: Salt Lake City, Utah, Utah Geological Association Publication 33 (in press). Paepke, B.E., 2001, Controls on channel organization in a glaciated basin in the Uinta Mountains, Utah [M.S. thesis]: Logan, Utah State University, 115 p. Paulsen, T., and Marshak, S., 1999, Origin of the Uinta Recess, Sevier foldthrust belt, Utah; Influence of basin architecture on fold-thrust belt

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geometry: Tectonophysics, v. 312, p. 203–216, doi: 10.1016/S00401951(99)00182-1. Powell, J.W., 1876, A report on the geology of the eastern portion of the Uinta Mountains and a region of country adjacent thereto: U.S. Geological Survey, 218 p. Pettengill, T., 1996, Lakes of the High Uintas: Sheep Creek, Carter Creek & Burnt Fork Drainages: Utah Division of Wildlife Resources Publication 96-17, 20 p. Reheis, M.C., Palmquist, R.C., Agard, S.S., Jaworowski, C., Mears, B., Jr., Madole, R.F., Nelson, A.R., and Osborn, G.D., 1991, Quaternary history of some southern and central Rocky Mountain basins: Bighorn basin, Green Mountain–Sweetwater River area, Laramie basin, Yampa River basin, northwestern Uinta basin, in Morrison, R.B., ed., Quaternary nonglacial geology: Conterminous U.S.: Boulder, Colorado, Geological Society of America, Geology of North America, v. K-2, p. 427–431. Richmond, G.M., 1965, Glaciation of the Rocky Mountains, in Wright, H.E., and Frey, D.G., The Quaternary of the United States: Princeton, New Jersey, Princeton University Press, p. 217–230. Richmond, G.M., 1986, Stratigraphy and correlation of glacial deposits of the Rocky Mountains, the Colorado Plateau and the ranges of the Great Basin: Quaternary Science Reviews, v. 5, p. 99–127, doi: 10.1016/02773791(86)90178-2. Ringen, B.H., 1984, Relationship of suspended sediment to streamflow in the Green River basin, Wyoming: U.S. Geological Survey Water-Resources Investigation Report 84-4026, 14 p. Ritzma, H. R., 1959, Geologic atlas of Utah, Daggett County: Utah Geological and Mineralogical Survey Bulletin 66, 116 p., scale 1:125,000. Ritzma, H.R., 1983. Igneous dikes of the eastern Uinta Mountains, Utah and Colorado: Utah Geological and Mineralogical Survey Bulletin 66, 111 p. Rowley, P.D., Hansen, W.R., Tweto, O., and Carrara, P.E., 1985. Geologic map of the Vernal 1° × 2° Quadrangle, Colorado, Utah, and Wyoming: U.S. Geological Survey Miscellaneous Investigations Series Map I-1526, scale 1:250,000. Schlenker, G.C., 1988, Glaciation and Quaternary geomorphology of the Blacks Fork drainage, High Uinta Mountains, Utah and Wyoming [M.S. thesis]: Salt Lake City, University of Utah, 87 p. Schoenfeld, M.J., 1969, Quaternary geology of the Burnt Fork area, Uinta Mountains, Summit County, Utah [M.S. thesis]: Laramie, University of Wyoming, 80 p. Sears, J.D., 1924, Relations of the Browns Park Formation and the Bishop conglomerate and their role in the origin of Green and Yampa Rivers: Geological Society of America Bulletin, v. 35, no. 2, p. 279–304. Shakun, J., 2003, Last Glacial Maximum equilibrium-line altitudes and paleoclimate, northeastern Utah [unpublished B.A. thesis]: Middlebury, Vermont, Middlebury College, 55 p. Sharp, W.D., Ludwig, K.R., Chadwick, O.A., Amundson, R., and Glaser, L.L., 2003, Dating fluvial terraces by 230Th/U on pedogenic carbonate, Wind River Basin, Wyoming: Quaternary Research, v. 59, p. 139–150, doi: 10.1016/S0033-5894(03)00003-6. Small, E.S., Anderson, R.S., Repka, J.L., and Finkel, R., 1997, Erosion rates of alpine bedrock summit surfaces deduced from in situ 10Be and 26Al: Earth and Planetary Science Letters, v. 150, p. 413–425, doi: 10.1016/S0012821X(97)00092-7. Smelser, M.G., 1997, Geomorphic adjustability of streams draining the Uinta Mountains [M.S. thesis]: Logan, Utah State University, 167 p. Stamp, M.E., 2000, Hydrologic and geomorphic effects of dams and water diversion on Lake Fork River and Rock Creek, Uinta Mountains, Utah [M.S. thesis]: Logan, Utah State University, 140 p. Stone, D.S., 1993, Tectonic evolution of the Uinta Mountains: Palinspastic restoration of a structural cross section along longitude 109°15′, Utah: Utah Geological and Mineral Survey Map 93-8, scale 1:96,000 and 1:192,000. Stone, J.O., 2000, Air pressure and cosmogenic isotope production: Journal of Geophysical Research, B, Solid Earth and Planets, v. 105, no. B10, p. 23,753–23,759, doi: 10.1029/2000JB900181. Trimble, K.W., 1924, Plan and profile of the Green River from Green River, Utah, to Green River, Wyoming: U.S. Geological Survey, 16 sheets (10 plans, 6 profiles), scale 1:31,680. Tweto, Ogden, 1976, Geologic map of the Craig 1° × 2° quadrangle, northwestern Colorado: U.S. Geological Survey Miscellaneous Investigations Series Map I-972, scale 1 250,000. Wahrhaftig, C., and Cox, A., 1959, Rock glaciers of the Alaska Range: Geological Society of America Bulletin, v. 70, no. 4, p. 383–436.

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Wallace, C.A., and Crittenden, M.D., 1969, The stratigraphy, depositional environment and correlation of the Precambrian Uinta Mountain Group, western Uinta Mountains, Utah, in Lindsey, J.B., ed., Geologic Guidebook of the Uinta Mountains: Intermountain Association of Geologists 16th Annual Field Conference, p. 127–142.

Westlund, T., 2005, Recessional moraines in the Yellowstone canyon, southern Uinta Mountains, Utah [unpublished B.A. thesis]: St. Peter, Minnesota, Gustavus Adolphus College, 26 p. Zimmer, T.M., 1996, Pedogenesis on quartzite-rich Pleistocene moraines, Smith’s Fork drainage, Uinta Mountains [M.S. thesis]: Logan, Utah State University, 118 p.

Printed in the USA

Geological Society of America Field Guide 6 2005

Geomorphology and rates of landscape change in the Fremont River drainage, northwestern Colorado Plateau David W. Marchetti Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112 USA John C. Dohrenwend Southwest Satellite Imaging, 223 South State Street, Teasdale, Utah 84773-0141 USA Thure E. Cerling Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112 USA

ABSTRACT The Fremont River drainage basin has a variety of geologic and geomorphic features that provide insight into the long-term landscape development of the catchment. Volcanic rocks that are ca. 26 to 4 Ma and are offset by Basin-and-Range style normal faulting underlie the western third of the drainage basin. Fish Lake Hightop and Boulder Mountain show evidence of Pleistocene glaciation. Recent mapping and surface exposure dating suggests that the glacial deposits around these two mountains were deposited during the last glacial maximum (LGM). Mass movement and fluvial deposits in the catchment are predominantly composed of volcanic boulders derived from the volcanic rocks atop Boulder and Thousand Lakes Mountains. Fremont River and tributary incision rates estimated from surface exposure dating of these deposits range from 0.20 to 0.43 m/k.y. Longer-term estimates of exhumation rates in the drainage basin based on emplacement depths of igneous rocks range from 0.10 to 0.38 km/m.y. Keywords: geomorphology, Colorado Plateau, landscape evolution, cosmogenic elements, glaciation, debris-flows. INTRODUCTION

mont River heads at Fish Lake, Utah, and runs east, crossing several structural features before joining with Muddy Creek to become the Dirty Devil River, a major tributary of the Colorado River. The river traverses ~100 km of the northwestern Colorado Plateau. The Fremont River drainage basin has a wide variety of rock types, amazingly displayed structural features, and an abundance of surficial deposits ranging in age from historic to Pliocene. The aim of this trip is to understand the effects of structure, process, and time on the development and evolution of the Fremont River basin. With that in mind, we pose the following questions as the intellectual focus of the trip: (1) What effect

“The Basin of the Colorado offers peculiar facilities for the study of the origin of topographic forms, and its marvelous sculpture has excited the interest of every observer.” —G.K. Gilbert (1877)

This three-day field trip examines the geology, geomorphology (marvelous sculpture), and rates of landscape change in the Fremont River drainage of south-central Utah (Figs. 1–3). The Fre-

Marchetti, D.W., Dohrenwend, J.C., and Cerling, T.E., 2005, Geomorphology and rates of landscape change in the Fremont River drainage, northwestern Colorado Plateau, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 79–100, doi: 10.1130/ 2005.fld006(04). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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Figure 1. Grayscale satellite image of the northwestern Colorado Plateau with some coverage of the Basin-and-Range and middle Rocky Mountains. The Fremont River basin is outlined in white. Some of the major features are labeled: GSL— Great Salt Lake; SLC—Salt Lake City; SRS—San Rafael Swell; HM—Henry Mountains; LSM—La Sal Mountains.

Geomorphology and rates of landscape change to I-15

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Figure 3. Grayscale satellite image of the Fremont River basin and the locations of stops for this field trip. Important geographical and structural features are labeled: GP—Geyser Peak; HHP—Hen Hole Peak; TLM—Thousand Lake Mountain; BM—Boulder Mountain.

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have structural features related to Laramide and Basin and Range deformation had on the Fremont River basin? (2) How have differences in rock strength affected landscape change? (3) What are the dominant geomorphic processes affecting landscape development in the headwaters and in the distal reaches? (4) What effects have Pleistocene glaciations had on the Fremont catchment? (5) How old are the various geomorphic deposits in the drainage basin? (6) Using dated features, what rates can be determined for river incision, overall landscape erosion, and Quaternary tectonic activity? While we certainly can’t answer all of these questions, we can provide some insight into several of them. Day 1 of this trip will start in the volcanic highlands of the Fish Lake, Awapa, and Aquarius Plateaus. We start with an overview of the northern part of the drainage basin, view the headwaters of the Fremont River, investigate the volcanic units on the Awapa Plateau, and end the day on a geomorphic surface offset by a Basin-and-Range normal fault. Day 2 of the trip will focus on mass movement deposits derived from the volcanic flows atop Boulder and Thousand Lakes Mountains, how these deposits have caused armoring and topographic inversion, and our cosmogenic dating of these deposits. We will start Day 2 in the Teasdale-Torrey area and end at the Capitol Reef National Park visitor center. On Day 3, we will look at more volcanic boulder–rich deposits along the Fremont River downstream of Capitol Reef, badlands formed in Mancos Shale near Caineville, Utah, and a suite of well-preserved fluvial terraces west of Hanksville, Utah. GEOLOGY OF THE FREMONT RIVER DRAINAGE Structure The Fremont River drainage basin lies within two major structural provinces of the western United States. The western end of the basin is within the Basin-and-Range–Colorado Plateau transition zone (TZ) and the eastern part of the basin lies within the Colorado Plateau (Fig. 1). The western edge of the Fremont River catchment (and western edge of the Awapa Plateau) is bounded by the Paunsagunt Fault, a N-S trending normal fault that is downthrown to the west. The next major fault to the east, the Thousand Lake fault, is considered the eastern boundary of the TZ and the beginning of the Colorado Plateau, as it is the easternmost expression of Basin-and-Range style normal faulting in the TZ (Fig. 3) (Wannamaker et al., 2001). The northern extension of the Thousand Lake fault is known as the Paradise fault zone (Nelson, 1989; Smith et al., 1963). Basin-and-Range extension began ca. 15 Ma and is still active (Stewart, 1978). The timing of last movement on Thousand Lake and Paradise faults is somewhat unclear. Rowley at al. (1979) suggest that most of the major displacements in the High Plateaus took place at ca. 7 Ma, whereas Nelson (1989) has field evidence for faulting younger than ca. 5 Ma in the Geyser Peak quadrangle, which includes part of the Thousand Lake and Paradise fault zones. Smith et al. (1963) and Anderson and Barnhard (1986) suggest that ~80–100 m of displacement has occurred on the Thousand Lake and Paradise

faults during the Pleistocene. The Utah Geological Survey online database of Quaternary faulting in Utah (www.geology.utah.gov/ maps/geohazmap/qfaults/imagemap2/index.html) suggests that the most recent event on the central Thousand Lake fault occurred during the mid- to late Quaternary (<750 ka). We will show evidence on Day 1 of the field trip that helps further constrain the timing of movement on the Thousand Lake fault. To the east of the Thousand Lake fault zone, the structural geology of the Fremont River drainage changes considerably. This part of the basin lies within the Colorado Plateau proper, where major Laramide-aged (90–50 Ma) structures are exposed. The most prominent Laramide feature that we will see in this field trip is the Miners Mountain anticline (Fig. 3). The Miners Mountain anticline is a W-NW–trending, asymmetrical, doubly plunging anticline that is the structural connection between the Waterpocket Fold (Circle Cliffs uplift) to the south and the San Rafael Swell to the north (Bump et al., 1997; Davis, 1999). Most references on the structural geology of Capitol Reef National Park hold that the steeply dipping sequence of rocks to the east and north of the visitor center (the “reef’ of Capitol Reef) are part of the Waterpocket Fold. However, Bump et al. (1997) and Davis (1999) argue that these features are instead part of the eastern limb of the Miners Mountain anticline. The anticline is bound on the southwest by the Miners Mountain fault and on the northwest by the Teasdale fault zone. The Teasdale fault zone has been interpreted as a high angle normal fault (Smith et al., 1963; Williams and Hackman, 1971), a series of tightly spaced NW-SE–trending folds (Billingsley et al., 1987), and as a reverse, left-handed, strike slip fault zone (Anderson and Barnhard, 1986; Davis, 1999). The Teasdale fault zone intersects the Thousand Lake fault at Red Gate, where the Fremont River crosses from the TZ into the Colorado Plateau (Fig. 3). It is possible that weaknesses in the local strata, caused by the Teasdale fault zone, allowed a subsequent proto–Fremont River to head or capture at this location, integrating the TZ portion of the Fremont River basin with the eastern Colorado Plateau portion. This intersection is also the likely reason Boulder and Thousand Lake Mountains are now isolated. The volcanic caps of Boulder and Thousand Lake Mountains are the same petrologically and in elevation, suggesting they were once connected. Igneous Rocks Three groups of igneous rocks are exposed in the Fremont River basin. Oligocene to Pliocene aged volcanic rocks from the Marysville volcanic field are exposed on the Fish Lake, Awapa, and Aquarius Plateaus in the western part of the basin (Fig. 3) (Anderson and Rowley, 1975; Rowley et al., 1994). The main eruptive phase of Marysville volcanics was between ca. 34 and 19 Ma and included large volumes of andesites, latites, and traychites, as well as volcanic breccias of those lithologies. Smaller episodic eruptions of rhyolitic tuffs and basaltic lava flows have occurred since then; some of the youngest eruptions are Pliocene in age (Rowley et al., 1979; Mattox, 1991).

Geomorphology and rates of landscape change Williams and Hackman (1971) describe the petrology and approximate ages for volcanics on the Fish Lake, Awapa, and Aquarius Plateaus. Mattox (2001) mapped the volcanics in the Moroni Peak quadrangle of the Awapa Plateau and provides revised petrologic interpretations and new numerical ages. The units, ages, and approximate thicknesses of volcanics in the Fremont River basin according to Williams and Hackman (1971) and Mattox (2001) are given in Table 1. The main volcanic unit that caps Boulder and Thousand Lakes Mountains, and the main lithology that makes up mass movement deposits derived from those mountains, is basaltic andesite (Tba) according to Williams and Hackman (1971) and potassium-rich mafic lava flows (Tpml) according to Mattox (2001). The youngest volcanic unit mapped by Mattox (2001) is a basalt lava flow that yielded a K-Ar age of 6.9 ± 0.3 Ma. This age is similar to a K-Ar age of 6.4 ± 0.4 Ma obtained by Delaney et al. (1986) for basalt from a vent on the top of Thousand Lake Mountain. The youngest ages of volcanic rocks in the drainage basin are from near Geyser Peak where Delaney et al. (1986) obtained a K-Ar age of 3.8 ± 0.2 Ma from a basalt flow near Hogan Pass and Nelson (1989) determined a K-Ar age of 3.95 ± 0.3 Ma for a trachybasalt from Geyser Peak. In the northern part of the Fremont River basin, several swarms of diabase and shonkinite dikes and occasional syenite plugs are exposed. Gartner (1986), Delaney et al. (1986), and Delaney and Gartner (1997) obtained K-Ar ages ranging from 3.4 ± 0.2 to 4.7 ± 0.3 Ma for the dikes and 4.6 ± 0.2 for a syenite sill. These ages are similar to the ages given above for basaltic flows near Geyser Peak and suggest a significant pulse of magmatic activity at that time (Nelson and Tingey, 1997). The dikes and plugs are shown on the maps of Williams and Hackman (1971) and Billingsley et al. (1987) and in Delaney and Gartner (1997). Stratigraphic relations suggest the dikes were emplaced between 0.5 and 1.5 km deep (Delaney and Gartner, 1997). In the southeastern part of the basin, intrusive siliceous rocks associated with the Henry Mountain laccoliths are exposed (Fig. 3). G.K. Gilbert first studied these intrusions in the late 1800s as the main focus of his famous monograph (Gilbert, 1877). The rocks that make up the laccolithic intrusions of the Henry Mountains are predominantly plagioclase-hornblende porphyries with much lesser amounts of syenite porphyry (Hunt et al., 1953; Nelson et al., 1992). 40Ar/39Ar and fission track dating suggests the laccoliths were emplaced between 20 and 31 Ma (Nelson, 1998; Sullivan, 1998). The timing of laccolithic intrusion in the Henry Mountains is therefore concurrent with the timing of the majority of volcanic activity of the Marysville volcanic field, located <50 km to the west. Jackson and Pollard (1988) estimated that the maximum emplacement depth of the Henry Mountain laccoliths was ~3–4 km (~1 kbar) based on structural and stratigraphic relationships. However, Nelson and Davidson (1998) suggest that ~4 km may be the minimum emplacement depth. They base their interpretation on the stability field of magmatic amphibole, which is unstable at pressures below ~1 kbar (depths less than ~4 km).

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TABLE 1. VOLCANIC UNITS ON THE AWAPA, AQUARIUS, AND FISH LAKE PLATEAUS Symbol

Name

Age (Ma)

Thickness (m)

Basalt lava flows Osiris Tuff K-rich mafic lava flows Latite lava flows Volcanic rocks of Langdon Mt.

6.9 ± 0.3 23 26–23 26–23 26–23

0–80 0–40 0–120 0–120 0–105

Pliocene Miocene Oligocene-Miocene Oligocene-Miocene Oligocene-Miocene

0–90 0–60 0–200 0–150 0–750

Mattox, 2001 Tb To Tpml Tl Tla

Williams and Hackman, 1971 QTb Tlo Tba Tl Tvb

Olivine basalt Tuff of Osiris Basaltic andesite Latite Volcanic breccia

Since they rarely see “breakdown” textures of magmatic amphiboles in Henry Mountain rocks, they conclude the pressure at the time of laccolithic emplacement was >1 kbar and the laccoliths were emplaced deeper than 4 km. Permian to Tertiary Stratigraphy Permian to Tertiary sedimentary rocks are well exposed in Fremont River basin east of the Thousand Lake fault zone. The most prominent formations, their thicknesses, and symbols (used from now on instead of their longer names) are given in Table 2. Descriptions and paleoenvironmental interpretations for all these units can be found in Smith et al. (1963), Williams and Hackman (1971), and Billingsley et al. (1987), as well as in most other books that discuss the general geology of the Colorado Plateau. Quaternary Stratigraphy A wide variety of Quaternary (to Pliocene?) surficial deposits are present in the Fremont River drainage. We discuss the most prominent and important of these under separate headings below. Many of the deposits are mapped, using varied classification schemes, in Flint and Denny (1958), Smith et al. (1963), Williams and Hackman (1971), and Billingsley et al. (1987). In most of those maps the deposits appear to be correctly delineated; however, we disagree with many of the unit assignments and implied genetic interpretations. Glacial Deposits Despite the high elevation of much of the Fremont River basin, only Boulder Mountain and Fish Lake Hightop have convincing evidence for past glacial activity. Erosional and depositional evidence for glaciation on Boulder Mountain is described by Dutton (1890), Gould (1939), Flint and Denny (1958), and Osborn and Bevis (2001). The “top” of Boulder Mountain is

D.W. Marchetti, J.C. Dohrenwend, and T.E. Cerling

84

TABLE 2. STRATIGRAPHIC COLUMN OF SEDIMENTARY ROCKS EXPOSED IN THE FREMONT RIVER DRAINAGE BASIN System

Formation

Thickness (m)

Symbol

Tertiary

Flagstaff Limestone

150

Tf

Cretaceous

Mesaverde Fm Mancos Shale Masuk Mb Emery Ss Mb Blue Gate Shale Mb Ferron Ss Mb Tununk Shale Mb Dakota Ss Cedar Mountain Fm

90–120

Kmv

200–230 90–120 370–460 60–120 165–220 0–45 0–50

Kmm Kme Kmb Kmf Kmt Kd Kcm

Morrison Fm Brushy Basin Mb Salt Wash Mb Summerville Fm Curtis Fm Entrada Fm Carmel Fm Navajo Ss

60–110 30–150 15–60 0-55 120–275 60–305 290–430

Jmb Jms Js Jcu Je Jc Jn

Triassic

Kayenta Fm Wingate Ss Chinle Fm Shinarump Mb Moenkopi Fm

110 110 150–210 0–30 240–305

T Rk TR w TRc TRcs TRm

Permian

Kaibab Limestone Cutler Group

0–60 240+

Pk Pc

Jurassic

Note: Modified from Billingsley et al., 1987. Fm—Formation; Mb— Member; Ss—sandstone.

relatively flat volcanic plateau (underlain by Tba/Tpml; Table 1), ~180 km2 in area with elevations ranging from 3320 to 3449 m. The top of the mountain is an unusual landscape, hosting large expanses of glacially sculpted, polished, grooved, or striated bedrock. During the last glacial maximum (LGM), Boulder Mountain supported an ice cap with outlet glaciers spilling off the main ice mass at several locations (Osborn and Bevis, 2001; Marchetti et al., 2005). The most complete study of the glacial deposits around the flanks of Boulder Mountain was done by Flint and Denny (1958). They mapped the majority of coarse boulder diamicts around the mountain and assigned them to one of three drift units they considered late Pinedale (MIS 2) to Bull Lake (now considered MIS 6, then considered MIS 4) in age based on relative morphology, clast weathering, and soil development. Subsequent work by Waitt (1997, 2000), Osborn and Bevis (2001), and Marchetti et al. (2005) concluded that the oldest drift unit (Carcass Creek drift) of Flint and Denny (1958) is from debris flow deposition rather than glaciation. Marchetti et al. (2005) further suggest that no convincing evidence for a glacial advance older than early Pinedale (MIS

2; LGM) has yet been found on Boulder Mountain. Marchetti et al. (2005) determined 3He exposure ages of ca. 25–15 ka for boulders on moraines in the Fish Creek drainage of Boulder Mountain. The lowest elevation reached by Pinedale age outlet glaciers from Boulder Mountain is estimated at ~2500 m (Waitt, 2000; Osborn and Bevis, 2001). Our surficial mapping in the Blind Lake and Government Point quadrangles, and field evidence reported by Waitt (2000), suggest that the Pinedale outlet glaciers from Boulder Mountain left only thin and discontinuous outwash trains. This contradicts the interpretations of Flint and Denny (1958), who mapped many of the downstream volcanic boulder–rich deposits as glacial outwash. Howard (1970), Anderson et al. (1996), and Repka et al. (1997) also concluded that many of the downstream fluvial terraces of the Fremont River are capped with outwash from various Boulder Mountain glacial advances. The glacial history of the Fish Lake Hightop has received considerably less study than Boulder Mountain. Based on moraine morphology and relative boulder weathering, Hardy and Muessig (1952) suggest that there were two glacial advances from the Fish Lake Hightop during the Pleistocene. Moraines associated with their oldest advance were designated as Wisconsin I and assigned an age of Bull Lake, while the younger moraines were called Wisconsin II and assigned an age of Pinedale. Osborn and Bevis (2001) further describe the two moraine sets and agree with Hardy and Muessig (1952) that the younger of the two moraines (Wisconsin II) is likely Pinedale in age. We have four unpublished 3He ages from the Tasha Creek moraines (Wisconsin I and II) that suggest that the two moraine sets mapped by Hardy and Muessig (1952) are early and late Pinedale in age rather than Bull Lake and Pinedale. Fluvial Deposits Deposits related to the Fremont River and tributaries are widespread throughout the basin, and many of them are mapped in Smith et al. (1963), Williams and Hackman (1971), and Billingsley et al. (1987). Most deposits associated with the mainstem Fremont River either contain or are dominated by volcanic clasts derived from the Marysville volcanic rocks exposed on the Fish Lake, Awapa, and Aquarius Plateaus in the western part of the basin. Where the Fremont River crosses the eastern limb of the Miners Mountain anticline (through Capitol Reef), there are well-formed and preserved strath terraces that are often rich in volcanic clasts. We use the term strath to describe a fluvially beveled bedrock surface and the term strath terrace to include the strath and the thin alluvium capping it (Bull, 1991; Pazzaglia et al., 1998). Thicker alluvial deposits that represent fluvial aggradation are called fill deposits. Where it is unclear if a terrace is strath or fill, we use the generic term “fluvial terrace” to refer to the deposits. In the Fremont drainage we only use the term fluvial terrace when we can clearly identify the drainage that the terrace was formed in; in most cases this means the terrace is still within a canyon or clearly related to a canyon. We constrain the usage this way to avoid confusion with landforms that we call “armored surfaces,”

Geomorphology and rates of landscape change which cannot be easily related to a particular drainage and are often topographically inverted features of great antiquity. We will further define these features below. Everitt et al. (1997) report that the alluvial fill of the modern Fremont River floodplain can be up to 30 m thick and that the thickness varies considerably along the course of the river. Mass Movement Deposits A variety of mass movement deposits are present throughout the Fremont River basin. Complex rotational slumps, translational landslides, massive debris flows, and smaller rock falls, topples, and avalanches have all been identified and mapped (Billingsley et al., 1987; Flint and Denny, 1958; Williams and Hackman, 1971; Smith et al., 1963). Most of the mappable mass movement deposits in the basin are derived from Boulder Mountain, Thousand Lake Mountain, or the northern outliers of Thousand Lake Mountain (Hen Hole and Geyser Peaks) (Fig. 3). Although there are mass movement deposits derived from the Fish Lake and Awapa Plateaus, there are considerably fewer of them, and they are not nearly as large as those from Boulder and Thousand Lake Mountains. Boulder and Thousand Lake Mountains are particularly susceptible to mass movement due to the presence of the Flagstaff Limestone (Tf) underlying the basaltic-andesites (Tba/Tpml; Table 1) that cap the mountaintop. The Paleocene-Eocene Flagstaff Limestone unit consists of limestones, weakly cemented tuffaceous sediments, thin sands, silts, and pebble conglomerates (Smith et al., 1963; Billingsley et al., 1987). The Flagstaff Limestone is similar to other Tertiary formations on the Colorado Plateau in that it is easily eroded and susceptible to mass movement (Godfrey, 1978). On Boulder and Thousand Lake Mountains, a large portion of the drainage off the mountaintop is through vertical fractures in the volcanic rock. This water emerges as springs at the contact with the Tf, causing sapping and over-steepening of the Tba/Tpml cliffs (Williams, 1984). Subsequent failure of these over-steepened cliffs leads to rock fall, rotational slumping, and landsliding; further mobilization of this detritus with the addition of water can trigger massive debris flows (Williams, 1984). The Awapa and Flagstaff Limestone Plateaus are underlain with volcanic units from summit to valley (Table 1). There is no “weak” layer for failure to occur on, so mass wasting deposits are less common. Rotational slumps and landslides can be seen on the north and east sides of Geyser and Hen Hole Peaks, all around Thousand Lake Mountain, and on most aspects of Boulder Mountain. Particularly spectacular rotational slumps occur on the north and west sides of Boulder Mountain at Government Point and Pine Creek Cove (Fig. 3). In summary, most of the coarse, volcanic, boulder-rich debris in the Fremont River catchment east of the Thousand Lake fault zone is from mass movement on Boulder or Thousand Lake Mountains. From our field observations, we suggest that most of this debris is first mobilized as rotational slumps or break-away toreva blocks from the steep volcanic cliffs (Tba/Tpml) ringing

85

the mountains. This detritus is then mobilized into massive debris flows either concurrent with the original failure, or later, most likely due to saturation during wet periods. These debris flows travel various distances down-valley, injecting occasional pulses of very coarse volcanic debris into the hill slope and fluvial systems. Piedmont Deposits Most piedmont deposits in the Fremont River drainage are located on the flanks of Boulder Mountain, Thousand Lake Mountain, and Hen Hole and Geyser Peaks. Deposits on these piedmonts are dominated by volcanic boulder clasts. The most common landscape element on these piedmonts is a volcanic boulder–rich debris flow to coarse alluvium capped pediment. There is considerable confusion surrounding the term pediment, and a consensus on the definition of this term is lacking (Cooke et al., 1993; Dohrenwend, 1994; Twidale, 1978). Here we use the term in the most general sense. We consider a pediment to be a beveled bedrock surface that slopes away from the upland. Our definition does not suggest any relation between pediments and base-level stability or past climatic conditions. A common landscape feature in the Fremont River drainage is a volcanic boulder–mantled pediment that has been incised and abandoned. We call these features “armored surfaces” because of the effect that the volcanic gravel cap has on the subsequent erosion rate of the surface. By determining the exposure ages of large boulders in armored surfaces, we can determine when these surfaces were last active and use the depth of exhumation around the surface to estimate local incision and overall landscape erosion rates. SURFACE EXPOSURE DATING At several of the stops on this trip we will discuss cosmogenic 3He exposure ages that we determined for boulder clasts exposed on the surfaces of various surficial deposits. The theory behind cosmogenic exposure age dating is complicated enough that we cannot treat it fully here. We direct interested readers to Gosse and Phillips (2002) as the most exhaustive review paper in the field. Additional, shorter reviews include Lal (1988, 1991), Bierman (1994), and Cerling and Craig (1994b). We used 3He concentrations, measured in pyroxene separates, to determine the emplacement age of volcanic boulders. We used a 3He production rate of 116 atoms g−1 yr−1 (Cerling and Craig, 1994a; Licciardi et al., 1999) and the scaling parameters of Lal (1991) to scale production rates to the latitude and altitude of each sample site. Further information on the corrections and analytical procedures we used to determine 3He concentrations and exposure ages is given in Marchetti and Cerling (2005) and Marchetti et al. (2005). Uncertainties reported with the exposure ages are 2σ. In determining the exposure ages, we assumed that our samples had negligible pre-exposure, zero erosion, and were not buried and then exhumed due to deposit erosion. The first of these assumptions is likely valid, as most of the deposits we sampled are very close to their source and are from debris flow

D.W. Marchetti, J.C. Dohrenwend, and T.E. Cerling

86

deposits, indicating relatively rapid exhumation, transportation, and deposition. In order to keep assumption one valid, we tried to avoid sampling lower or inset deposits that could easily collect clasts eroded from higher surfaces. We also resisted the temptation to sample any of the impressive strath terraces along the mainstem Fremont River, except where they were clearly capped by massive debris flow deposits (e.g., Stop 2-5). The second and third assumptions are almost certainly false as all subaerially exposed rocks and deposits undergo some erosion. To try to account for the effects of boulder and deposit erosion, we collected multiple samples from each surface and recorded the height of each sample relative to the deposit surface. By plotting the exposure age of each clast against its height above the deposit tread, we can qualitatively assess the effect of boulder and deposit erosion. On older surfaces (older than ca. 300 ka), where boulder and deposit erosion should control the exposure ages, one would expect to find a positive relationship between exposure ages and boulder height. The highest boulder above the surface would be the oldest, and the oldest age would be the closest to the “true” age of the deposit. This is often not the case and suggests that there are other variables controlling the exposure age distributions. Possibilities include differences in boulder erodibility, localized shielding effects, or possible preexposure. Since boulder and deposit erosion are the main sources of uncertainty in our exposure ages, and they cause ages to be too young, we interpret the oldest exposure age from each deposit as a minimum age for the deposit. Any incision or erosion rate determined from a minimum age would be a maximum rate. In determining the incision rates, we included an estimate of each deposit thickness, typically 2–10 m, but did not include an estimate for the present thickness of alluvial fill in the Fremont or tributary drainages, which can be considerable (Everitt et al., 1997). Not including this estimate would underestimate the incision depth in many locations and make the rates slightly too low. DAY 1. SALT LAKE CITY TO FREMONT JUNCTION TO TORREY, UTAH The road log starts at the Fremont Junction exit of Interstate 70 (Fig. 2). Note: UTM coordinates given for numbered stops are referenced to NAD 27, zone 12. The U.S. Geological Survey 7.5′ quadrangle for each stop is given after these coordinates. Many directions and modifiers are abbreviated to conserve space (Rt. = route, CR = county road, NW = northwest, etc.). Compass directions are azimuths corrected to true north with a 13°E declination. Cumulative mi (km) 0.0

(0.0)

Description At the stop sign at the end of the I-70 exit, turn right then right again onto Rt. 72. Drive parallel to interstate. Several volcanic boulder armored surfaces are in this area. Bedrock is Kmb (see symbols in Table 2).

1.9

(3.1)

2.8

(4.5)

6.8

(10.9)

7.4

(11.9)

9.1

(14.6)

12.8

(20.6)

17.4

(28.0)

17.8

(28.6)

18.1

(29.1)

Road veers left. At junction with Rt. 76, continue S on Rt. 72. Road continues up Post Hollow toward pass between Thousand Lake Mountain and Fish Lake Plateau. Cliffs are Kmm (Table 2) in Paradise fault zone. Entering Signboard Flat; Tertiary volcanics capping cliffs to the right. Entering Sage Hole, a relatively undissected perched alluvial filled fault bound valley. Entering Fish Lake National Forest, axis of Last Chance Creek; cliffs on the right are Kmm. Floor of Paradise Valley, a fault bounded internally drained depression on E flank of Fish Lake Plateau. View to the S of Thousand Lake Mountain, Hen Hole, and Geyser Peaks. View of Henry Mountains to the SE; entering Fremont River basin. Willow Basin sign; pull off to left. Stop 1-1.

Stop 1-1. Overview of Northern Capitol Reef and Fish Lake Plateau 0458722E 4270874N, Geyser Peak Quadrangle The objective of this overlook stop is to view many of the geomorphic features in the northern half of the Fremont River drainage and to get a feel for the faulted volcanic terrain of the Fish Lake Plateau. On a clear day, looking generally E-SE, you can see across the entire northern half of the Colorado Plateau from this point. The La Sal Mountains are at ~105°, the Abajo Mountains are at ~125°, and the Henry Mountains are at ~140°. North and South Caineville Mesas are visible near the base of the Henry Mountains. The Fremont River bisects these two mesas. To the south of this stop is Geyser Peak, which has been mapped as undifferentiated basaltic andesites and latites by Williams and Hackman (1971) and as trachybasalts and shoshonite lava flows by Nelson (1989). A K-Ar age of 3.95 ± 0.3 Ma was obtained by Nelson (1989) on trachybasalts near the crest of Geyser Peak. This age matches a K-Ar age of 3.8 ± 02 obtained from one of the many diabase dikes exposed to the east of here, in the northern part of Capitol Reef National Park (Gartner, 1986). On the eastern slope of Geyser Peak is a rotational slump. Nelson (1989) calls this feature the Solomon Basin landslide and suggests that it is Pleistocene in age. This style of mass wasting is similar to what occurs on the volcanic capped mountains to the south of here (Hen Hole Peak, Thousand Lakes Mountain, and Boulder Mountain). To the E-SE, in the middle distance, you can see several flat-topped armored surfaces. These surfaces are volcanic-boulder debris flow–capped pediments that have been inverted due to boulder armoring, differential erosion, and incision. The relief around these surfaces is on the order of 100–200 m. These and similar types of surfaces will be the

Geomorphology and rates of landscape change focus of much of Day 2. Understanding the formation and ages of these armored surfaces is key to understanding the long-term landscape development of the Fremont River basin. Cumulative mi (km) 18.4

(29.6)

22.2

(35.7)

25.0 26.4

(40.2) (42.5)

29.0

(46.7)

29.8 33.1 33.8

(47.9) (53.3) (54.4)

35.4

(57.0)

Description Kiosk on left, views of Fish Lake Plateau to the right, descending down to Fremont River valley. Great view of Thousand Lake fault zone ahead and to the left, Fish Lake Hightop to the right, area dominated by Oligocene to Pliocene age volcanics modified by normal faulting. Wayne County line. Turn right onto Rt. 25 toward Mill Meadow Reservoir and Fish Lake. Sevier County line; good views to the left of high Fremont River terraces. Road crosses Fremont River. Exposure of bedded volcaniclastics on the left. Scenic view; pullout on the right; drive on to next scenic view 1.6 mi ahead. Pull into scenic view pullout. Stop 1-2.

Stop 1-2. Upper Fremont River 0447110E 4270132N, Forsyth Reservoir Quadrangle The purpose of this stop is to view the headwaters of the Fremont River and discuss 3He exposure ages of clasts from a high fluvial terrace. Looking down the drainage, you can see

87

Thousand Lake Mountain straight ahead and Boulder Mountain off to the right in the far distance. Exposures of Tf are visible on the left side of Thousand Lake Mountain just below the summit. On the right side of the canyon you can see a series of very high, flat topped, fluvial terraces (Fig. 4). These terraces are capped with coarse boulder diamicts and are isolated by incised tributaries draining Mytoge Mountain, located to the south. Determining the thickness of the deposits capping these terraces is difficult because the local bedrock is andesitic volcanic breccia and the terraces are capped with andesitic volcanic boulder diamict. We determined 3He exposure ages for three boulders from the largest terrace, located just downstream of Ivy Canyon (Fig. 4). The ages range from 607 ± 25 ka to 508 ± 21 ka, and the oldest exposure age is from the highest and most unweathered clast. Using the oldest exposure age and a measured tread to modern floodplain distance of ~155 m, we estimate an average incision rate of 0.26 m/k.y. for this stretch of the Fremont River. Remember, exposure ages are considered minimum ages and therefore rates are maximum rates. Upstream of here the Fremont River emerges from Johnson Valley Reservoir and leaves the Fish Lake graben (Fig. 4). The Fremont River appears to be structurally controlled (subsequent) through this canyon, roughly following a NW-SE fault trace on the NE flank of the Fish Lake graben. Several other similar faults control the tributaries to the NE. The relationship between Fish Lake, the Fish Lake graben, and Fremont River integration is unclear. Hardey and Muessig (1952) suggest that Fish Lake once drained to the SW, out onto the Awapa Plateau. Webber (2003) rejects that hypothesis due to lack of field evidence and suggests that perhaps Fish Lake was internally drained after graben formation (initiated ca. 6–7 Ma) and then later captured by an ancestral Fremont River.

FLH JVR (2688 m)

Fremont River

M ou nt ai n M yt og e

Fi (2 sh 69 La 5 ke m )

stop 1-2

N ~ 3 km

Figure 4. 10 m digital elevation model of the Fish Lake graben and upper Fremont River (illumination at 315° and 30°). Stop 1-2 is shown with a white star. The Ivy Canyon terrace is outlined with a solid white line. Other possibly correlative terraces are outlined with dashed white lines. FLH—Fish Lake Hightop; JVR—Johnson Valley Reservoir.

D.W. Marchetti, J.C. Dohrenwend, and T.E. Cerling

88 Cumulative mi (km) 35.4 44.4

48.5 52.9 54.2 55.1 57.3

57.7 58.9 61.1 61.9

62.7

Description

(57.0)

Turn around and head back down Rt. 25 toward Rt. 72. (71.4) Junction with Rt. 72; turn right. Geyser Peak ahead; to the left, Jc exposed at base. Road descends into Rabbit Valley. (78.0) Road crosses Fremont River. (85.1) Sharp turn to the right, heading toward Loa. (87.2) Rt. 72 ends; at junction with Rt. 24, turn left and head through Loa. (88.7) As Rt. 24 veers to the left, take the right spur and head south on South Main Street. (92.2) As the road veers to the left, turn right onto unmarked gravel road. This is the main access road to the Awapa Plateau. (92.8) Core stone-like outcrops of Miocene Osiris Tuff (To/Tlo; Table 1). (94.8) Coming out of Osiris and into Tba/Tpml. (98.3) Road drops into Big Hollow; Awapa Plateau rises to the W. (99.6) Ford Big Hollow. Big Hollow is one of the most deeply incised drainages on all of the Awapa Plateau. Quaternary olivine basalt (Qtb/Tb) on slope above. (100.9) Pull off road for Stop 1-3.

Stop 1-3. Volcanic Rocks South of Loa, Utah 0446270E 4239750N, Bicknell Quadrangle At this stop we will look at some of the volcanic rocks exposed on the Awapa Plateau and get good views of the important geomorphic features of the western part of the drainage basin (Fig. 3). From here you can see the Awapa Plateau at 180°–315°, the west side of Boulder Mountain from 120°– 160°, Red Gate, where the Fremont River crosses the Thousand Lake fault and the Colorado Plateau beginning at ~90°, Thousand Lake Mountain at ~50°, the Fremont River through Grass Valley at ~0°, and Fish Lake Hightop, Fish Lake graben, and Mytoge Mountain at ~340°. The petrology and ages of the volcanic rocks exposed on the Awapa Plateau are given in Table 1. The unit exposed here is Tl according to both Williams and Hackman (1971) and Mattox (2001). Note the relative lack of dissection across much of the Awapa Plateau to the west. Most of the rocks here are late Oligocene to early Miocene in age and are only slightly dissected. Cumulative mi (km)

Description

62.7 (100.9) Proceed S on gravel road. 63.2 (101.7) At junction, turn left and head E, traveling down dip on volcanics. 65.2 (104.9) Stop 1-4.

Stop 1-4. View of Red Gate 0448790E 4240355N, Bicknell Quadrangle This short stop is designed to give us better views of the Thousand Lake fault zone, Red Gate, and the W side of Boulder Mountain. Looking ~60°–70° there are good exposures of much of the TR sequence on the flank of Thousand Lake Mountain. At the base of Thousand Lake Mountain in the same direction you can see a downthrown and tilted block of volcanic rock capping an exposure of Tf. To the south of this block is a beheaded alluvial fan, which is our next stop. At ~90° there are great views down the axis of Red Gate. You can see the offset of TR to J units on the left and right sides of Red Gate caused by the reverse motion of the Teasdale Fault. At this spot the strike of the Teasdale Fault zone projects directly at us. Between ~140° and 170° are views of the faulted and up-thrown west side of Boulder Mountain. A variety of mass wasting and glacial deposits are exposed on the west side of Boulder Mountain. Preliminary 3He ages of boulders from some of these deposits range from ca. 130 to 20 ka. These features are only slightly dissected and do not appear to be offset by faulting. Cumulative mi (km)

Description

65.2 (104.9) Continue E toward Fremont River. 66.6 (107.2) Merge with other gravel road—bear left, continue heading N. 67.0 (107.8) Road becomes paved and crosses Fremont River. 68.6 (110.4) Intersection of 100S and 400W; proceed N on 400W. 68.8 (110.7) Junction of Rt. 24 and 400W; turn right and head E through Bicknell. 69.0 (111.0) Utah State liquor agency on left: only liquor store in Wayne County. Welcome to Utah! 69.9 (112.5) Turn left onto Sunglow campground road. Road crosses N flank of abandoned alluvial fan. 70.3 (113.1) Pull off on the right, where the dirt road heads E uphill and park. Stop 1-5. Stop 1-5. Offset Bicknell Fan and Thousand Lake Fault 0454243E 4243094N, Bicknell Quadrangle Walk up the two-track road ~500 m, until the road meets a volcanic boulder–strewn hill on your left. Climb up the hill for Stop 1-5. This stop is in the Thousand Lake fault zone (Figs. 3 and 5). The unit capping this ridge seems to be andesitic volcanic breccia, suggesting it is Tvb or Tla (Table 1). Underneath the volcanic rocks (exposed just north of here) is an outcrop of bright white Tf, which occurs as a friable ash deposit here. Beneath that, and to the NE, is a large, buff to yellow-orange block that is also likely Tf but could be Kmv (Table 2).

Geomorphology and rates of landscape change

89

Bicknell Sand

Figure 5. Aerial photograph of the old Bicknell piedmont (OBP) surface near the town of Bicknell, Utah. The head of the offset fan is outlined with several arrows.

h Was stop 1-5

OBP

N ~ 0.5 km

Sloping to the south of this tilted block is a volcanic boulder–covered surface. Given the large size of some of the clasts (2–3 m b-axis) and the discrete grouping of the big boulders, this surface looks like a debris flow fan. The volcanic clasts on this surface are dominantly basaltic andesite (Tba/Tpml; Table 1) and are not the same as the volcanics capping the ridge we are on. The only upslope source of these volcanics is the outcrop of basaltic-andesite that caps Thousand Lake Mountain. Looking to the NE, it becomes clear that the fan surface has been beheaded and abandoned. The fan head is truncated right at the easternmost segment of the Thousand Lake fault (Fig. 5), suggesting that it was offset and abandoned due to faulting. We determined the 3He exposure ages of four large volcanic (Tba/Tpml) boulders from the abandoned fan surface. The exposure ages (and measured boulder heights) are 154 ± 8 ka (1.8 m), 83 ± 6 ka (1.1 m), 182 ± 8 ka (1.3 m), and 213 ± 9 ka (1.5 m). The exposure ages have considerable scatter and suggest that either we sampled clasts from separate debris-flows deposits having different ages or that boulder erosion has strongly affected some of the exposure ages. If the youngest exposure age of 83 ± 6 ka is accurate and not erroneously young, then the surface was offset in the past ~83 k.y. However, if the young ages are too young due to boulder erosion and the oldest exposure age (213 ± 9 ka) is the best estimate (still a minimum age) of the age of the surface, then the offset occurred sometime since deposition of that clast.

Looking directly south, you can see the effect that the Thousand Lake fault has had on the Fremont River (Fig. 6). The marshy area just to the west of the fault zone is known as Bicknell Bottoms and is an area of considerable aggradation as the Fremont River deposits sediment in order to maintain its course across the up-thrown block at Red Gate. The major convexity in the Fremont River long profile is due to both the Thousand Lake fault flattening the Fremont River profile at Bicknell Bottoms and the resistant Permian units (Pk and Pc) exposed downstream in the Miners Mountain anticline, steepening the profile in the Fremont River gorge (Fig. 6). Cumulative mi (km)

Description

70.3 (113.1) Turn around and head back to Rt. 24. 70.7 (116.2) Junction with Rt. 24; turn left and head toward Torrey. 72.2 (116.2) Off of abandoned fan, within Thousand Lake fault zone. 72.8 (117.1) Highway veers to the left and enters G.K. Gilbert’s “Red Gate.” 74.8 (120.4) Road climbs up onto volcanic boulder– capped Teasdale Bench. Road cuts show gravel-pediment contact. 75.1 (120.8) Turn off to Teasdale on the right, continue on Rt. 24 toward Torrey. End of road log for Day 1.

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90

DAY 2. ARMORED SURFACES AROUND TEASDALE AND CAPITOL REEF NATIONAL PARK Cumulative mi (km) 0.0

(0.0)

0.7

(1.1)

1.3

(2.1)

1.8

(2.9)

2.2

Description Start Day 2 road log at junction of Rt. 24 and Great Western Trail (GWT) just E of Torrey, Utah. Turn onto GWT and head S. Boulder Mountain fills skyline ahead; Miners Mountain is to the left with the Henry Mountains beyond. Small ridge of TRm, dropping into Fremont River Canyon. Multiple levels of Fremont River terraces to the right, State Stone Corp. Up out of canyon onto low boulder-capped terrace. To SE, view of cockscomb—upturned Jn in Teasdale fault zone. Junction with CR-3262 (Teasdale/Grover road). Turn right, heading toward Teasdale. Turn right onto unmarked dirt road; small quarry on the left. Road forks; take left fork. Pull into turnoff on right and drive down to turnaround. Park here. Stops 2-1 (here) and 2-2 (up boulder-capped hill to W).

(3.5)

2.6

(4.2)

3.2

(5.2)

3.5 3.8

(5.6) (6.1)

3.9

(6.3)

3000

E

2800

W

2600

Elevation (m)

2400 2200

Stop 2-1. Fremont River Canyon Overlook 0462608E 4236555N, Torrey Quadrangle The purpose of this short stop is to get an overview of the geomorphic features in the Teasdale area (Fig. 7). Looking ~0°– 45°, you can see the SE flank of Thousand Lake Mountain. In this view are at least five major boulder-armored surfaces, most of which are completely abandoned and inverted. The Fremont River cuts though a small canyon at ~110°. Notice several levels of strath terraces with volcanic boulders. At ~160° is the trace of the Teasdale fault zone, where you can see a large upturned block of Jn known as the cockscomb. The wooded slopes to the S of the cockscomb are all capped with volcanic boulder debris from the NE flank of Boulder Mountain. Many of these armored surfaces are now completely inverted as well. We will walk ~1 km to the W-NW up onto the large armored surface for stop 2-2. Stop 2-2. Inverted Teasdale Surface 0462100E 4236724N, Torrey Quadrangle We call this feature the TEA surface due to its closeness to the town of Teasdale (Fig. 7). This feature is a great example of topographic inversion as it was once the valley floor and is now a high mesa. The bedrock under the deposit capping the TEA surface is TRm. The stratigraphy of the deposit and size of the clasts exposed on the surface suggest that this feature is capped with massive debris flow deposits. We determined 3He exposure ages for three boulders from the TEA surface. The ages range from 556 ± 25 ka to 475 ± 22 ka and the exposure ages show no correlation with boulder height. The range of exposure ages is most likely due to variable boulder erosion sometime after the clasts were deposited, but could also be due to hillslope exposure prior to deposition. Relief around the TEA surface ranges from ~50 m to 120 m and the Fremont River is incised ~110 m below the NE edge of the surface. Using the oldest boulder exposure age of 556 ka, we estimate an incision rate of ~0.20 m/k.y. for this reach of the Fremont River. Cumulative mi (km)

2000 1800

3.9

(6.3)

4.6

(7.4)

5.0

(8.1)

5.8

(9.3)

7.5

(12.1)

1600 1400 1200

v.e. = 65x

1000 0

20

40

60

80

100

120

140

160

Distance upstream from confluence with Muddy Creek (km) Figure 6. Valley longitudinal profile of the Fremont River constructed from 10 m digital elevation models. The profile starts at the Johnson Valley reservoir (Fig. 4). The approximate location and relative motion of the Thousand Lake fault is shown with a dashed line. The axis of the Miners Mountain anticline is shown with an arrow.

Description Turn around and backtrack to paved CR3262. Junction with CR-3262; turn right. Driving on a low, armored surface; slope is roughly parallel to that of stop 2-2. Great view of TEA surface on the right. Younger debris flow fans on left from Boulder Mountain drainages. Driving on mix of older boulder armored surfaces. In Teasdale fault zone. Note deformation bands in Jn and TRw on right side of road. Road makes abrupt right turn into downtown Teasdale.

Geomorphology and rates of landscape change 7.7

(12.4)

7.9

(12.7)

8.1

(13.0)

8.3 8.5

(13.4) (13.7)

Turn left just before the church, heading W up Bullberry Creek canyon. Outcrops of Jn capped by Jc, still in Teasdale fault zone. On older part of Bullberry Creek, debris flow fan surface. Teasdale rodeo grounds on left. As road veers to the left, turn right into turnout by small perched holding pond. Walk up onto ridge by pond and over to the top of the boulder strewn ridge for stop 2-3.

91

capped by Jc; at ~0°, in the distance, is Thousand Lake Mountain, and in the middle distance are exposures of brilliant orange TRw dipping toward the S. At 90°–135°, the present drainage of Bullberry Creek debouches onto a fan at the town of Teasdale, and at ~180° to 200° is the canyon of Bullberry Creek with exposures of Jc on top of Jn. Bullberry Creek heads on boulder-covered slopes on northernmost Boulder Mountain. The drainage is just to the west of a huge rotational slump that, morphologically, looks to be quite young (Fig. 7). The streamlined ridges in the center of the slump and the presence of deformed tree trunks suggest that it may have remobilized very recently. The debris flow deposits along Bullberry Creek can be traced on the aerial photo (Fig. 7) for several kilometers upstream to the northern slope of Boulder Mountain. Presently, Bullberry Creek has incised ~5 m along the southern edge of the deposits, and a smaller unnamed tributary has incised ~5–10 m on the northern edge of the deposits. The Bullberry

Stop 2-3. Bullberry Creek Debris Flows 04569442E 4237298N Torrey The purpose of this stop is to examine the Bullberry Creek debris flow deposit (Fig. 7). At ~270° are exposures of dipping Jn

Holt Draw Bench

Red Gate TB2

Torrey Bench Torrey

Fremont River TB1 BR

BCF 2-3

2-4 Teasdale

TEA 2-2 2-1

toe of remobilized slump N ~1.5 km Figure 7. Aerial photograph of the important geomorphic and geographic features in the Teasdale-Torrey area. Numbered stops are shown with white stars. The Fremont River flows from W to E. BCF—Bullberry Creek fan; BR—Black Ridge; TEA—Teasdale surface; TB1—older Teasdale Bench; TB2—younger Teasdale Bench.

D.W. Marchetti, J.C. Dohrenwend, and T.E. Cerling

92

Creek deposits appear to grade into the lower Teasdale Bench (TB2; Fig. 7) deposits that are the focus of the next stop. We determined 3He exposure ages for three boulders from the BCF surface. The exposure ages (and boulder heights) are 224 ± 10 ka (1.1 m), 105 ± 5 ka (1.3 m), and 200 ± 9 ka (0.9 m). All the samples were collected from the same general area, but the boulder yielding the youngest exposure age is located slightly upstream and to the west of the other two samples. We suggest that the two groups of exposure age represent boulders emplaced during two different debris flow events. However, given that we have only three ages, we realize that this suggestion is somewhat tenuous. The younger exposure age could be erroneously young due to deposit or boulder erosion. The two older exposure ages are similar to ages from the lower TB surface (TB2) that we will see at the next stop. The morphology of the BCF surface is remarkably similar to the much older TB1 surface (again, next stop), and we suggest that BCF is a good analogue for what that much older surface used to look like (Fig. 7).

source of previously exposed clasts is the boulder deposits atop Black Ridge, which is located ~300 m above and to the west of the BCF and TB surfaces (Fig. 7). The exposure ages from the TB2 surface are 184 ± 8 ka (1.3 m) and 258 ± 16 ka (0.9 m). The ages on the TB2 surface are roughly similar to the two older ages from the BCF surface (224–200 ka) and suggest that the two surfaces may have been active at the same time. Cumulative mi (km) 9.9

(15.9)

10.8

(17.4)

12.3

(19.8)

12.8

(20.6)

14.5 15.1 17.1 19.1 23.7

(23.3) (24.3) (27.5) (30.8) (38.2)

Stop 2-4. Teasdale Bench

25.2

(40.6)

0458360E 4237946N, Torrey Quadrangle From the pullout on CR-3263, walk up to the ridge crest to the south, crossing the barbed wire fence at a low point. At this stop we will investigate the Teasdale Bench (TB) surfaces. We informally identify two different TB surfaces here; TB1 is the older, higher surface we are on, and TB2 is the lower much more extensive debris flow–capped fan surface that spreads out to the north all the way to the Fremont River (Fig. 7). The TB2 surface sits ~4–8 m below the TB1 surface remnant. Both surfaces slope N-NE toward the Fremont River. The huge clast sizes (~2 m baxis), poor sorting, and discrete grouping of boulders suggest that these deposits are remnants of massive debris flows. As noted in the previous stop, the southwestern taper, morphology, and slope of the TB1 surface are very similar to the Bullberry Creek deposit where it emerges from Bullberry Creek canyon. We sampled three boulders from the TB1 surface and two boulders from the TB2 surface. The ages (and boulder heights) from the TB1 surface are 426 ± 18 ka (1.2 m), 451 ± 19 ka (1.4 m), and 602 ± 25 ka (1.3 m). We suspect that the ca. 600 ka age is older because of pre-exposure. This is based on the ca. 556 ka age of the TEA surface, which is considerably higher and the possibility that clasts from higher surfaces in the area could have been incorporated into the TB deposits. The most likely

25.7

(41.4)

Cumulative mi (km) 8.5 9.4

(13.7) (15.1)

9.9

(15.9)

Description Turn around and head back to CR-3262. At junction with CR-3262, turn left, heading N. Good views of the slope of TEA surface on the right. Road cut through older Teasdale Bench (TB) deposit; pull off to the right just past road cut. Walk back uphill, crossing fence, for stop 2-4.

Description Continue north on CR-3262. Good views of TB2 fan surface to the right; outcrops of TRm, TRms, TRc, and TRw dead ahead. Junction with Rt. 24, turn right heading toward Torrey. Heading E across TB2 surface, note boulders, surface relief, relative lack of dissection. Crossing Fremont River, climbing up onto Torrey Bench. On Torrey Bench, occasional road cuts and outcrops show pediment surface. GWT trail turn off on the right. Junction with Rt. 12; continue on Rt. 24. Crossing Sand Creek. Entering Capitol Reef National Park. As road descends, great views of Johnson Mesa debris flow deposit (stop 2-5) dead ahead. Junction with Capitol Reef scenic road. Turn right and proceed past visitor center. Make sharp right turn onto dirt service road; proceed ~75 m to parking area on the left. We’ll walk up the road for stop 2-5.

Stop 2-5. Johnson Mesa Hike up the service road. Stop at the third major switchback near pipes and buried tanks. Stop 2-5a. Johnson Mesa Road Cut; 0477556E 4237328N, Twin Rocks Quadrangle The purpose of this first stop is to view the internal stratigraphy of the Johnson Mesa deposit (Figs. 8 and 9). At ~270°–90° are exposures of the TR to J stratigraphic sections exposed on the eastern limb of the Miners Mountain anticline. The road cut here exposes fluvial sands and gravels at the base of the Johnson Mesa deposit. Three informal units are seen between here and ~50 m further up the service road: Unit 1: Unit 1 is at the base of the exposure here and is ~0.6 m thick. Thinly bedded pebbles to fine sands dominate the unit. Alternating white and dark layers reflect the relative amounts of the basaltic andesite component. In one area, the beds are tilted to the NE and appear to be prograding foresets.

Geomorphology and rates of landscape change

o

o

111 20’

111 15’

Johnson Mesa Terrace 1780-1710 m

Outline of debris flow deposit Boulder sample Desert pavement sample

o

38 17’

93

Figure 8. Digital orthophoto and sketch map of Johnson Mesa terrace, Carcass Creek terrace, Fremont gorge, and the Fremont River.

Carcass Creek Terrace ~2050 m

Fre mon

t Ri

ver

N

o

38 16’

km Fremont River gorge Carcass Creek

Bedded layers show occasional cut and fill features. Unit 1 has a gradational contact with Unit 2, and is interpreted as a lowenergy fluvial/floodplain deposit. Unit 2: Unit 2 is ~1 m thick, clast supported, and is dominated by subangular to rounded basaltic andesite clasts with some sandy horizons. Basaltic andesite clasts range in size from pebbles to ~1 m in diameter (b-axis). Clasts have moderate imbrication and weakly coarsen upward. The contact with Unit 3 is distinct. Unit 2 is interpreted as a coarse high-energy fluvial deposit. Unit 3: Unit 3 extends from the top of Unit 2 to the surface of the deposit and is ~6–8 m thick. Unit 3 is unsorted, matrixsupported, and dominated by coarse angular to subangular basaltic andesite clasts ranging in size from pebbles to ~2 m in diameter (b-axis). The matrix is composed of muds to fine sands and is white to gray in color. This unit is interpreted as a massive debris flow deposit. Continue up the service road another ~50 m. Notice the huge basaltic andesite clast on your left that has exposed dimensions of 12 × 6 × 4 m and is matrix supported. The only conceivable way to move this boulder ~20 km from its source on Boulder Mountain would have been by entrainment in either a high-density slurry flow or ice. Continue another ~100 m up the service road. On your left is an inset terrace cut into the Johnson Mesa deposit. This terrace was likely cut by Sulfur Creek and is capped with a much finer grained deposit that has considerably fewer basaltic

0

1

2

Johnson Mesa

Figure 9. Photograph of the Johnson Mesa terrace. Boulder Mountain is in the distant background and Fremont gorge is in the left background. U.S. Highway 24 in the center foreground for scale.

andesite clasts. Follow the service road along the ditch for another 200–300 m to stop 2-5b. Stop 2-5b. Johnson Mesa Surface; 0477809E 4237054N, Twin Rocks Quadrangle At this stop, we are on the surface of the Johnson Mesa deposit (Figs. 8 and 9). At ~90° is the canyon of the Fremont

D.W. Marchetti, J.C. Dohrenwend, and T.E. Cerling

94

River where it crosses the eastern limb of the Miners Mountain anticline. At ~240° is Fremont gorge which is the canyon of the Fremont River across the axis of Miners Mountain anticline (Fig. 8). The walls of Fremont gorge expose Pk on top of Pc. TRm is exposed on the crest of the fold and on most of the slopes seen to the west. We determined 3He exposure ages for three large boulders from the Johnson Mesa deposit (Fig. 10) (Marchetti and Cerling, 2005). The ages group tightly together and suggest that the debris flow capping the surface was deposited at ca. 193 ± 10 ka (minimum age). A similar debris flow deposit called the Carcass Creek Terrace is located ~8 km upstream of Johnson Mesa (Fig. 8). Exposure ages of two boulders from the CCT surface are within uncertainties of the ages obtained for the Johnson Mesa surface (Fig. 10). Using the boulder exposure ages, Marchetti and Cerling (2005) determined maximum incision rates of 0.43 and 0.40 m/k.y. for the Johnson Mesa and Carcass Creek reaches of the Fremont River. In addition to the exposure ages, we obtained three 230Th/ U ages for one well-developed carbonate rind that formed on the underside of a boulder on Johnson Mesa (Fig. 11). Warren Sharp analyzed the samples at the Berkley Geochronology Center (see Sharp et al., 2003, for methods). The ages preserve stratigraphic order along the growth axis of the crust and give a relatively constant rate of carbonate precipitation. The projected age of “first” carbonate precipitation for the crust is ca. 180 ka (Fig. 11). This age estimate is likely too low as there is some time lag between deposit emplacement and significant pedogenic carbonate precipitation. Sharp et al. (2003) suggest that this time lag is ~5 ± 5 k.y. for carbonate rinds formed on glacial outwash from the Wind River Range, Wyoming. Adding ~5 k.y. to the projected age yields an age of ca. 185 ka, which is similar to, but younger than the 3He exposure ages. However, we should note that accurate 230Th/U rind dating of Quaternary deposits requires considerably more rind samples and age transects before the age of the landform can be reliably estimated (Sharp et al., 2003). Cumulative mi (km) 25.7

(41.4)

26.4

(42.5)

Description Turn around and turn left onto Capitol Reef scenic road heading back toward visitor center. Junction with Rt. 24; turn left heading toward Torrey. End of road log for Day 2.

DAY 3. TORREY TO HANKSVILLE ALONG THE FREMONT RIVER Cumulative mi (km) 0.0

(0.0)

Description Start log for Day 3 at Capitol Reef National Park Visitor Center.

0.1

(0.2)

1.2

(1.9)

3.0

(4.8)

3.8 6.8

(6.1) (10.9)

7.8

(12.6)

9.3

(15.0)

9.8

(15.8)

9.9

(15.9)

Milepost 80—highway turns into Fremont River canyon and follows course of river for next 40 mi. Milepost 81—Petroglyph turnoff on the left. Road level is at base of TRw. For next 6 mi, the highway will proceed upsection through the Glen Canyon group (but downstream through the canyon through Capitol Reef). Low boulder-capped strath on the right. Watch for numerous, discontinuous volcanic boulder–capped straths on both sides of canyon for the next 4 mi. These are most numerous and best developed where cut on TRk. Remains of 1997 rockfall of Jn on the right. Artificial cut and diversion of Fremont River—abandoned meander on the right, several straths and strath deposits in artificial rincon. Highway enters South Desert, the valley of Deep Creek—a strike valley paralleling Capitol Reef and underlain by Je—also entering zone of complete topographic inversion. Leave Capitol Reef National Park—junction with Rt. 1670 to Notom and Bullfrog. Turn right and head uphill. Turn right onto dirt road and veer left at fork; drive out to large turnaround. Stop 3-1.

Stop 3-1. South Desert Overlook 0488582E 4236486N, Fruita Quadrangle Great views all around of steeply dipping sequences of Jn to Jms exposed on the eastern limbs of the Miners Mountain anticline (NW to SW) and Circle Cliffs uplift (SW to SE). Structurally discordant beveled surfaces are also exposed in every direction; these are either armored surfaces (higher inverted features) or strath terraces (lower, encanyoned). We are standing on a paleo-surface carved on Je, with cliffs of Js immediately around us to the E-SE. To the N-NW, Deep Creek is incised into a strike valley on the eastern limb of the Miners Mountain anticline in an area known as South Desert (Fig. 12). Patches of Deep Creek terraces are located to the E of the main drainage axis. Deep Creek and similar drainages formed on dipping strata of varying competence often display what G.K. Gilbert (1877) termed monoclinal shifting. At ~50° to 60° is the highest armored surface in the area, which we call the Blue Desert Bench. The surface slopes to the W-SW, roughly the same direction as the slope of the modern Fremont River. 3He exposure ages (and boulder heights) of two boulders from the Blue Desert Bench are 570 ± 37 ka (0.5 m) and 466 ± 62 ka (0.6 m). The oldest boulder age of 570 ± 37 ka is estimated as the minimum age for the deposit since the clasts were small and somewhat weathered. Using that age and an

Geomorphology and rates of landscape change

150

230

Devils Hole interglacial

marine interglacial MIS 5e

100

95

180

CCT boulders

Th / U age (ka)

JM boulders

200

average growth rate: 0.16 mm/kyr

160 140 120 100 80

crosses are 2σ

JM desert pavements

3

He exposure age (ka)

250

50

60 0

4

8

12

16

20

Distance from clast (mm)

0

Figure 10. 3He exposure ages and uncertainties of boulders and desert pavements from the Johnson Mesa and Carcass Creek terraces. The approximate duration of the penultimate interglacial period according to the Devils Hole δ18O record (Winograd et al., 1997) is shown with a light gray band. The approximate duration of the penultimate interglacial period according to marine δ18O records (Imbrie et al., 1984) is shown with a darker gray band. The weighted mean exposure age for each surface is shown with a dashed line. JM—Johnson Mesa; CCT—Carcass Creek terrace.

Figure 11. 230Th/U ages of pedogenic carbonate sampled from a carbonate rind from Johnson Mesa. The samples were collected from a transect perpendicular to the carbonate growth laminations. The oldest carbonate is located at the clast/rind contact (0 on x–axis).

De ep Cr

North Blue Flats

ee

Red Desert

Blue Gate

k BDB

S.Caineville Mesa

stop 3-1

N ~ 3 km

Figure 12. Grayscale satellite image of the Fremont River canyon between Capitol Reef National Park and Blue Gate. The Capitol Reef National Park visitor center is shown with a white square. Stop 3-1 is shown with a white star. North Blue Flats lies in a small, asymmetrical structural basin with Kmt at its center, whereas Red Desert lies in a small asymmetrical structural dome exposing Je at its center. The prominent ridge between Red Desert and Blue Gate is Kmf capping Kmt and is known as the Caineville Reef. South Caineville Mesa is capped with Kme atop Kmb. BDB—Blue Desert Bench (see text).

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96

incision depth of ~166 m we estimate a maximum incision rate of 0.29 m/k.y. for this reach of the Fremont River. A similar armored surface is located ~16 km to the NW of the Blue Desert Bench along Hartnet Draw and Deep Creek. Exposure ages of two 0.7-m-high clasts from that surface are ca. 900 ka and yield a similar maximum incision rate of 0.30 m/k.y. for that northern Fremont River tributary. Cumulative mi (km) 9.9 10.5 10.7

(15.9) (16.9) (17.2)

12.7

(20.4)

13.3 13.6

(21.4) (21.9)

14.2

(22.8)

15.2 16.0

(24.5) (25.7)

17.0 20.1

(27.4) (32.3)

20.4

(32.8)

21.2

(34.1)

23.0

(37.0)

Description Turn around heading back downhill to Rt. 24. Turn right onto Rt. 24. Highway enters canyon cut into Js, Jms, and Jmb. Boulder-capped strath cut into Jmb (Blue Desert Bench) ahead on skyline. This surface is the highest landscape element in this area. Milepost 91—during the August 2003 flash flood, the water of the Fremont River was ~1 ft deep on the road at this point. Highway crosses Pleasant Creek. View down the Fremont River—boulder gravel-capped straths flank both sides of the canyon. Notom cutoff road on right. Approximate location of contact between Jmb and Kmt. Highway climbs through Kmt. Crest of saddle in Kmt—highway descends back to Fremont River canyon. Two prominent strath mesas lie ahead in middle distance. North Caineville Mesa forms skyline beyond. Highway crosses Fremont River. Highway leaves strike valley and curves E through Caineville Reef where Caineville Wash cuts through Kmf cap. Highway crosses Caineville Wash and enters G.K. Gilbert’s “Blue Gate.” Highway proceeds through Caineville badlands for next 7 mi. South Caineville Mesa on the right; North Caineville Mesa on the left. When you see the famous “800 ft” cement truck on the left, slow down! Pull off on left side of road and pass through gate to Stop 3-2.

Stop 3-2. Caineville Badlands 0501615E 4244254N, Town Point Quadrangle The Caineville badlands are the most extensive and well developed badlands on the Colorado Plateau. Measurements from a systematic analysis of Landsat 7 satellite data show that these badlands cover a contiguous area of almost 40,000 acres (~162 km2; ~63 mi2). This area is nearly 50% of the combined area of all badland terrain in the Colorado Plateau, yet represents less

than 0.1% of the total area of the plateau. Most of the other welldeveloped badlands on the plateau occur as small areas (<2.5 km2; <1 mi2) that are widely scattered across the region. Thus, these badlands represent an unusual and distinctive landscape. The Caineville badlands are developed in the Blue Gate Member of the Mancos Shale (Kmb). In the Caineville area, the Kmb unit is essentially flat-lying, nearly 450 m thick, and highly susceptible to erosion. This shale is overlain by an ~100m-thick, erosionally resistant sandstone (Kme) that caps Factory Butte as well as both North and South Caineville Mesas. Distinguishing features of the Caineville badlands are (1) closely spaced, sharp ridge crests separated by deep, narrow valleys; (2) high drainage densities; (3) steep slopes. Slopes ranging between 35° and 40° occupy a large proportion of the landscape, and in areas immediately adjacent to North and South Caineville Mesas, nearly all of the terrain is >30°; and (4) extremely low vegetation densities. Vegetation coverage is typically less than 1% in most areas and approaches 0% on most slopes. Low mean annual precipitation (~13 cm; ~5 in) and the chemical and physical properties of the Blue Gate Shale combine to make these badlands about as close to an absolute desert as can be found in North America. Large areas of the Caineville badlands have been significantly impacted by off-highway–vehicle (OHV) activity over the past ~25 yr. Comparative geomorphic analysis of the surficial characteristics of undisturbed versus heavily disturbed hillslopes in these badlands suggests that OHV use is accelerating erosion. The absence of soil crust, rills and small gullies, and colluvium on hillslopes that have been heavily disturbed by OHV activity demonstrates that at least 7–13 cm of additional erosion has occurred on these hillslopes as compared to natural, undisturbed slopes. This amount of additional erosion is equivalent to a soil loss of one million pounds (~4.5 × 105 kg; ~500 tons) per hillslope acre (~4050 m2). Cumulative mi (km) 23.0

(37.0)

24.7 24.8

(39.7) (39.9)

28.4

(45.7)

28.9

(46.5)

29.2

(47.0)

Description Return to vehicles; continue on Rt. 24 heading E. Mesa Market. Birdie Wash—at this point, we are crossing the axis of the Henry Basin. Bedrock at this point is nearly flat-lying. For the next 2.5 mi, the highway crosses a broad flat transport surface cut on Kmb. Slow down! Turn off onto gravel road on the right (road to Factory Butte is to the left). If you miss this turn, another right turn comes up in about 0.1 mi and will get you to the same place. Continue uphill to top of “airport” strath terraces. Road forks; take right fork and proceed to terrace edge near “airport” control buildings. Stop 3-3.

Geomorphology and rates of landscape change Stop 3-3. Old Hanksville Airport Terrace 0509366E 4245611N Town Point The purpose of this stop is to view the geology of the Henry Mountain basin and a series of well-preserved strath terraces of the Fremont River (Fig. 13). There are good views of the northern Henry Mountains to the south. To the W-SW is the valley of the Fremont River we just drove through with North and South Caineville Mesas (Kme over Kmb). Factory Butte (also Kme over Kmb) and the eastern edge of the San Rafael Swell are located to the north. We are standing on a major strath terrace of the Fremont River cut into Kmf; several other terrace levels are visible to the west. Howard (1970) first studied these terraces and their relationship to the inception of badland formation in the Caineville area. He correlated these terraces with the glacial stages proposed by Flint and Denny (1958) for Boulder Mountain. Repka et al. (1997) identified four levels of Fremont terraces in the Caineville area that they labeled FR-0 (modern floodplain) to FR-4 (Fig. 13). We are standing on their FR-3 terrace looking west toward patches of FR-3 and the multiple levels of FR-2. Anderson et al. (1996) and Repka et al. (1997) used cosmogenic 10 Be and 26Al measured in amalgamated surface quartzite clast samples, along with similar measurements on buried clasts from depth profiles, to estimate the ages of the most extensive of these terraces. The depth sampling and amalgamation is needed to account for the significant cosmogenic pre-exposure that fluvial clasts can acquire during exhumation, hillslope residence, and transport. In Table 3 we report individual clast and terrace exposure ages from Repka et al. (1997) and 3He exposure ages we determined on volcanic clasts from two of the terraces.

97

To the south and north of this stop are exposures of intrusive igneous rocks that were emplaced at quantifiable depths. Using the radiometric ages and emplacement depths of these rock bodies, we can estimate long-term exhumation rates for the Fremont River catchment. The Henry Mountains laccoliths have been dated at ca. 20–31 Ma (Nelson, 1998; Sullivan, 1998) and are thought to have been emplaced between ~3– 4 km (Jackson and Pollard, 1988) and >4 km deep (Nelson and Davidson, 1998). The top of the intrusions now stand out with ~2 km of positive relief, suggesting a total exhumation range of ~3–6 km. Using that depth range and the age range given above yields long-term average exhumation rates of ~0.10–0.30 km/ m.y. A similar calculation can be performed using data from diabase dikes exposed in the northern part of Capitol Reef National Park. Based on stratigraphic relationships, Delaney and Gartner (1997) suggest that the dikes were emplaced between 0.5 and 1.5 km deep. Using that estimate and an average K-Ar age of ca. 4 Ma (Delaney et al., 1986; Gartner, 1986; Delaney and Gartner, 1997) yields average exhumation rates of ~0.13–0.38 km/m.y. In Table 4 we show a summary of the incision rates determined in this guidebook and the exhumation rates determined above. Cumulative mi (km) 29.2 29.9 30.6

(47.0) (48.1) (49.2)

36.7

(59.1)

Description Return to vehicles and head back to Rt. 24. Turn right at junction with Rt. 24. For next 3.5 mi the highway is on a transport surface cut on Kmt. Highway climbs up onto a low strath terrace of the Fremont River.

Rt. 24

N 1 km

stop 3-3

FR-3

FR-2c FR-1

Fremont River FR-3

Figure 13. Aerial photograph of the Fremont River east of Caineville, Utah. Stop 3-3 is shown with a white star. Several strath terrace levels are shown as mapped by Repka et al. (1997).

D.W. Marchetti, J.C. Dohrenwend, and T.E. Cerling

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TABLE 3. EXPOSURE AGE DATA FOR THE FREMONT RIVER TERRACES NEAR CAINEVILLE, UTAH Single clast 10Be age range, no correction (ka)

10 Be age of terrace from clast amalgamation and depth profiles (ka)

FR-0

6–20

N.D.

N.D.

FR-1

24–125

N.D.

N.D.

FR-2a FR-2b FR-2c

64 60–69 56–130

N.D. N.D. 60 ± 9

N.D. N.D. N.D.

FR-3

84–202

102 ± 16

161 ± 9–208 ± 11

FR-4

N.D.

151 ± 24

288 ± 22

Repka et al., 1997

Repka et al., 1997

Terrace

Reference

3

He age range (ka) of volcanic clasts, no correction (± 2σ) (FR-3; n = 3)

Note: n—number; N.D.—not determined.

39.8 41.2

(64.0) (66.3)

Highway crosses Fremont River. Junction Rt. 24 and Rt. 95; turn left onto Rt. 95 heading N. End of road log.

CONCLUDING REMARKS 1. The western third of the Fremont River drainage basin is considerably different than the eastern two-thirds as it is underlain by a sequence of Miocene to Pliocene aged volcanic and volcaniclastic rocks. The age of these volcanic units and their exact stratigraphic sequence is not well understood. Additional large-scale mapping and improved geochronology would provide a better framework for developing hypotheses regarding the inception and development of the upper Fremont River drainage and its relations Colorado River integration. 2. The Thousand Lake fault is an important structural and geomorphic feature in the drainage basin. The fault has a significant effect on the long profile of the Fremont River and is the easternmost expression of Basin and Range extension in the basin. Cosmogenic ages from an offset debris flow fan near Bicknell, Utah, suggest that the fault may have been active in the past ~213 k.y. 3. Pleistocene glaciation was restricted to Boulder Mountain and Fish Lake Hightop. No conclusive evidence for glacial advances older than Pindedale (LGM, MIS 2) has yet been found on either mountain. In both locations, the Pinedale glaciers left only slight amounts of outwash. 4. Volcanic boulder–rich mass movement deposits from Geyser, Hen Hole, Thousand Lake, and Boulder Mountains are common throughout the drainage basin. These deposits cover many of the slopes surrounding the mountains and occasionally traveled many km farther down valley as massive debris flows. Over time, the relative toughness of the boulders in these deposits

can cause the surfaces they mantle to be armored, incised around, and topographically inverted. 5. Rock strength exerts a fundamental control on the geomorphology of the Fremont River drainage basin. The cuesta form landscape in the eastern two-thirds of the basin clearly displays the effect of relative rock strength on hillslope morphology and cliff recession. Additionally, the extremely resistant rock units exposed by the Miners Mountain anticline have formed a major kick-point in the Fremont River long profile. 6. Bedrock incision rate estimates from exposure-dated clasts on armored surfaces and river terraces range from 0.2 to 0.8 m/k.y. Long-term exhumation rate estimates based on the emplacement depths of dated igneous rock bodies range from 0.1 to 0.3 km/m.y. Although comparing rates of geomorphic processes determined over significantly different time intervals is not always appropriate (Gardner et al., 1987), the k.y. and m.y. rates for the Fremont River basin are both relatively fast and indicate an active and changing landscape. Much of this incision was likely driven by base-level lowering since Colorado River integration occurred in the latest Miocene to early Pliocene (Lucchitta, 1972). ACKNOWLEDGMENTS We thank Kip Solomon and Alan Rigby for help with He analyses. Richard Waitt, Jamie McCaughey, Ben Passey, Elliot Lips, Ian Schofield, Chuck Bailey, Cassie Fenton, Scott Hynek, Will Gallin, and Suzanne Bethers all helped with fieldwork at various times. Capitol Reef National Park provided logistical support for both our research and this field trip. The Grand Staircase–Escalante National Monument and U.S. Geological Survey educational mapping component of the National Cooperative Geologic Mapping Program (EDMAP) helped fund this research. Joel Pederson provided helpful reviews that significantly improved this chapter.

Geomorphology and rates of landscape change TABLE 4. SUMMARY OF INCISION AND EXHUMATION RATE ESTIMATES FOR THE FREMONT RIVER DRAINAGE BASIN Location

Fremont River Ivy Canyon terrace (Stop 1-2) Fremont River TEA surface (Stop 2-2) Fremont River Johnson Mesa terrace (Stop 2-5) Fremont River Blue Desert Bench (Stop 3-1) Tributary drainage—Deep Creek Hartnet surface (Stop 3-1) Fremont River Airport terraces (age data from Repka et al., 1997) (Stop 3-3) Diabase dikes (emplaced 0.5–1.5 km deep at ca. 4 Ma) Henry Mountains (emplaced 3–6 km deep at 20–31 Ma)

Incision/exhumation rate estimate (m/k.y.) 0.26

0.20

0.43

0.29

0.30

0.68–0.87

0.13–0.38 0.10–0.30

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Delaney, P.T., and Gartner, A.E., 1997, Physical processes of shallow mafic dike emplacement near the San Rafael Swell, Utah: Geological Society of America Bulletin, v. 109, p. 1177–1192, doi: 10.1130/00167606(1997)109<1177:PPOSMD>2.3.CO;2. Delaney, P.T., Pollard, D.D., Ziony, J.I., and McKee, E.H., 1986, Field relations between dikes and joints: Emplacement processes and paleostress analysis: Journal of Geophysical Research, v. 91, no. B5, p. 4920–4938. Dohrenwend, J.C., 1994, Pediments in arid environments, in Abrahams, A.D., and Parsons, A.J., eds., Geomorphology of desert environments: London, Chapman & Hall, p. 321–353. Dutton, C.E., 1890, Report on the geology of the high plateaus of Utah: U.S. Geographical and Geological Survey of the Rocky Mountains Region (Powell), v. xxxii, 307 p. Everitt, B.L., Godfrey, A.E., Anderson, R.S., and Howard, A.D., 1997, Quaternary geology and geomorphology, northern Henry Mountains region: BYU Geology Studies, v. 42, p. 373–404. Flint, R.F., and Denny, C.S., 1958, Quaternary geology of Boulder Mountain Aquarius Plateau, Utah: Geological Survey Bulletin, v. 1061-D, p. 103–164. Gardner, T.W., Jorgensen, D.W., Shuman, C., and Lemiex, C.R., 1987, Geomorphic and tectonic process rates: Effects of measured time interval: Geology, v. 15, p. 259–261, doi: 10.1130/0091-7613(1987)15<259:GATPRE>2.0.CO;2. Gartner, A.E., 1986, Geometry, emplacement history, petrography, and chemistry of a basaltic intrusive complex, San Rafael and Capitol Reef areas, Utah: U.S. Geological Survey Open File Report 86-81, 112 p. Gilbert, G.K., 1877, Geology of the Henry Mountains: Geographical and Geological Survey of the Rocky Mountain Region: U.S. Government Printing Office. Godfrey, A.E., 1978, Land surface instability on the Wasatch Plateau, central Utah: Utah Geology, v. 5, no. 2, p. 131–141. Gosse, J.C., and Phillips, F.M., 2001, Terrestrial in situ cosmogenic nuclides: theory and application: Quaternary Science Reviews, v. 20, p. 1475–1560, doi: 10.1016/S0277-3791(00)00171-2. Gould, L.M., 1939, Glacial geology of Boulder Mountain, Utah: Geological Society of America Bulletin, v. 50, p. 1371–1380. Hardy, C.T., and Muessig, S., 1952, Glaciation and drainage changes in the Fish Lake Plateau: Geological Society of America Bulletin, v. 63, p. 1109–1116. Howard, A.D., 1970, A study of process and history in desert landforms near the Henry Mountains, Utah [Ph.D. thesis]: Johns Hopkins University, 197 p. Hunt, C.B., Averitt, P., and Miller, R.L., 1953, Geology and geography of the Henry Mountains region Utah: U.S. Geological Survey Professional Paper 228, 234 p. Imbrie, J., Hays, J.D., McIntyre, A., Mix, A.C., Morley, J.J., Pisias, N.G., Prell, W.L., and Shackleton, N.G., 1984, The orbital theory of Pleistocene climate: Support from a revised chronology of the marine d18O record, in Berger, A., Imbrie, J., Hays, J., Kukla, G., and Saltzman, B., eds., Milankovich and climate: Boston, Reidel, p. 269–305. Jackson, M.D., and Pollard, D.D., 1988, The laccolith-stock controversy: New results from the southern Henry Mountains, Utah: Geological Society of America Bulletin, v. 100, p. 117–139, doi: 10.1130/00167606(1988)100<0117:TLSCNR>2.3.CO;2. Lal, D., 1988, In situ-produced cosmogenic isotopes in terrestrial rocks: Annual Review of Earth and Planetary Sciences, v. 16, p. 355–388. Lal, D., 1991, Cosmic ray labeling of erosion surfaces: in-situ nuclide production rates and erosion models: Earth and Planetary Science Letters, v. 104, p. 424–439, doi: 10.1016/0012-821X(91)90220-C. Licciardi, J.M., Kurz, M.D., Clark, P.U., and Brook, E.J., 1999, Calibration of cosmogenic 3He production rates from Holocene lava flows in Oregon, USA, and effects of the Earth’s magnetic field: Earth and Planetary Science Letters, v. 172, p. 261–271, doi: 10.1016/S0012-821X(99)00204-6. Lucchitta, I., 1972, Early history of the Colorado River in the Basin-and-Range province: Geological Society of America Bulletin, v. 83, p. 1933–1947. Marchetti, D.W., and Cerling, T.E., 2005, Cosmogenic 3He exposure ages of Pleistocene debris flows and desert pavements in Capitol Reef National Park, Utah: Geomorphology, v. 67, no. 3-4, p. 423–435, doi: 10.1016/ j.geomorph.2004.11.004. Marchetti, D.W., Cerling, T.E., and Lips, E.W., 2005, A glacial chronology for the Fish Creek drainage of Boulder Mountain, Utah: Quaternary Research, in press. Mattox, S.R., 1991, Petrology, age, geochemistry, and correlation of the Tertiary volcanic rocks of the Awapa Plateau, Garfield, Piute, and Wayne counties, Utah: Utah Geological Survey Miscellaneous Publication 91-5, 46 p.

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Mattox, S.R., 2001, Geologic map of the Moroni Peak quadrangle, Wayne County, Utah: Utah Geologic Survey Miscellaneous Publication 01-5, scale 1:24,000. Nelson, S.T., 1989, Geologic map of the Geyser Peak quadrangle, Wayne and Sevier counties, Utah: Utah Geological and Mineral Survey Map 114, scale 1:24,000. Nelson, S.T., 1998, Reevaluation of the central Colorado Plateau laccoliths in the light of new age determinations, in Friedman, J.D., and Huffman, A.C., eds., Laccolith complexes of southeastern Utah: Time of emplacement and tectonic setting—workshop proceedings: U.S. Geological Survey Bulletin 2158, p. 37–39. Nelson, S.T., and Davidson, J.P., 1998, The petrogenesis of the Colorado Plateau laccoliths and their relationship to regional magmatism, in Friedman, J.D., and Huffman, A.C., eds., Laccolith complexes of southeastern Utah: time of emplacement and tectonic setting—workshop proceedings: U.S. Geological Survey Bulletin 2158, p. 85–100. Nelson, S.T., and Tingey, D.G., 1997, Time-transgressive and extension-related basaltic volcanism in southwest Utah and vicinity: Geological Society of America Bulletin, v. 109, no. 10, p. 1249–1265, doi: 10.1130/00167606(1997)109<1249:TTAERB>2.3.CO;2. Nelson, S.T., Davidson, J.P., and Sullivan, K.R., 1992, New age determinations of central Colorado Plateau laccoliths, Utah: Recognizing disturbed K-Ar systematics and re-evaluating tectonomagmatic relationships: Geological Society of America Bulletin, v. 104, p. 1547–1560, doi: 10.1130/00167606(1992)104<1547:NADOCC>2.3.CO;2. Osborn, G., and Bevis, K., 2001, Glaciation in the Great Basin of the Western United States: Quaternary Science Reviews, v. 20, p. 1377–1410, doi: 10.1016/S0277-3791(01)00002-6. Pazzaglia, F.J., Gardner, T.W., and Merrits, D.J., 1998, Bedrock fluvial incision and longitudinal profile development over geologic time scales determined by fluvial terraces, in Tinkler, K., and Wohl, E., eds., River over rock: Fluvial processes in bedrock channels: Washington D.C., American Geophysical Union, 323 p. Repka, J.L., Anderson, R.S., and Finkel, R.C., 1997, Cosmogenic dating of fluvial terraces, Fremont River, Utah: Earth and Planetary Science Letters, v. 152, p. 59–73, doi: 10.1016/S0012-821X(97)00149-0. Rowley, P.D., Steven, T.A., Anderson, J.J., and Cunningham, C.G., 1979, Cenozoic stratigraphic and structural framework of southwestern Utah: U.S. Geological Survey Professional Paper 1149, 22 p. Rowley, P.D., Mehnert, H.H., Naeser, C.W., Snee, L.W., Cunningham, C.G., Steven, T.A., Anderson, J.J., Sable, E.G., and Anderson, R.E., 1994, Isotopic ages and stratigraphy of Cenozoic rocks of the Marysvale volcanic field and adjacent areas, west-central Utah: U.S. Geological Survey Bulletin 2071, 35 p.

Sharp, W.D., Ludwig, K.R., Chadwick, O.A., Amundson, R., and Glaser, L.L., 2003, Dating fluvial terraces by 230Th/U on pedogenic carbonate, Wind River Basin, Wyoming: Quaternary Research, v. 59, p. 139–150, doi: 10.1016/S0033-5894(03)00003-6. Smith, J.F., Huff, L.C., Hinrichs, E.N., and Luedke, R.G., 1963, Geology of the Capitol Reef area, Wayne and Garfield Counties, Utah: Geological Survey Professional Paper 363, 102 p. Stewart, J.H., 1978, Basin and Range structure in western North America: A review, in Smith, R.B., and Eaton, G.P., Cenozoic tectonics and regional geophysics of the western Cordillera: Geological Society of America Memoir 152, p. 1–31. Sullivan, K.R., 1998, Isotopic ages of igneous intrusions in southeastern Utah, in Friedman, J.D., and Huffman, A.C., eds., Laccolith complexes of southeastern Utah: Time of emplacement and tectonic setting—workshop proceedings: U.S. Geological Survey Bulletin 2158, p. 33–35. Twidale, C.R., 1978, On the origin of pediments in different structural settings: American Journal of Science, v. 278, p. 1138–1176. Waitt, R.B., 1997, Huge Pleistocene (Pliocene?) debris avalanches from Aquarius Plateau through Waterpocket Fold Monocline, Utah: Geological Society of America Abstracts with Programs, v. 29, no. 6, p. 410. Waitt, R.B., 2000, Road Log, in Waitt, R.B., Marchetti, D.W., Cerling, T. E., Kreutzer, L., and Anderson, A., eds., Red Gate to Blue Gate: Field guide to Friends of the Pleistocene: Rocky Mountain section, 44th annual reunion, p. 23–50. Wannamaker, P.E., Bartley, J.M., Sheehan, A.F., Jones, C.H., Lowry, A.R., Dumitru, T.A., Ehlers, T.A., Holbrook, W.S., Farmer, G.L., Unsworth, M.J., Hall, D.B., Chapman, D.S., Okaya, D.A., John, B.E., and Wolfe, J.A., 2001, Great Basin–Colorado Plateau transition in central Utah: An interface between active extension and stable interior, in Erskine, M.C., Faulds, J.E., Bartley, J.M., and Rowley, P.D., eds., The geologic transition, High Plateaus to Great Basin—A symposium and field guide: Salt Lake City, Utah Geological Association, p. 1–38. Webber, C.E., 2003, Structural and fluvial history of the Fish Lake Basin, High Plateaus, central Utah [B.S. thesis]: Williamsburg, Virginia, College of William and Mary, 76 p. Williams, P.L., and Hackman, R.J., 1971, Geology, structure, and uranium deposits of the Salina quadrangle, Utah: U.S. Geological Survey Miscellaneous Geologic Investigations Map I-591. Williams, V.S., 1984, Pedimentation versus debris-flow origin of plateau-side desert terraces in southern Utah: Journal of Geology, v. 92, p. 457–468. Winograd, I.J., Landwehr, J.M., Ludwig, K.R., Coplen, T.B., and Riggs, A.C., 1997, Duration and structure of the past four interglaciations: Quaternary Research, v. 48, p. 141–154, doi: 10.1006/qres.1997.1918.

Printed in the USA

Geological Society of America Field Guide 6 2005

Late Cretaceous stratigraphy, depositional environments, and macrovertebrate paleontology of the Kaiparowits Plateau, Grand Staircase–Escalante National Monument, Utah Alan L. Titus John D. Powell Grand Staircase–Escalante National Monument, Kanab, Utah 84741, USA Eric M. Roberts University of Witwatersrand, Johannesburg, South Africa Scott D. Sampson Utah Museum of Natural History, Salt Lake City, Utah 84112-0050, USA Stonnie L. Pollock Fasken Oil and Ranch, Ltd., Midland, Texas 79701-5116, USA James I. Kirkland Utah Geological Survey, Salt Lake City, Utah 84114-6100, USA L. Barry Albright University of North Florida, Jacksonville, Florida 32224, USA

ABSTRACT The Kaiparowits Basin, located mostly within Grand Staircase–Escalante National Monument, preserves an outstanding record of Late Cretaceous sedimentation in a foreland basin setting. Hosted in these rocks is one of the most continuous and complete records of this period’s ecosystems known from any one geographic area in the world. Recent work in the basin has emphasized macrovertebrate remains and documented many new sites of high scientific value. Recent stratigraphic studies have further refined our knowledge of the depositional systems and chronostratigraphic relationships. Provided is an overview of some of these recent advances, along with the necessary background to provide context. Keywords: Late Cretaceous, paleontology, Kaiparowits, vertebrate, stratigraphy, Utah.

Titus, A.L., Powell, J.D., Roberts, E.M., Sampson, S.D., Pollock, S.L., Kirkland, J.I., and Albright, L.B., 2005, Late Cretaceous stratigraphy, depositional environments, and macrovertebrate paleontology of the Kaiparowits Plateau, Grand Staircase–Escalante National Monument, Utah, Utah, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 101–128, doi: 10.1130/2005.fld006(05). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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Utah contains one of the best, most continuous records of Mesozoic terrestrial stratigraphy and paleontology in the world. Much of the Cretaceous part of this record has only been documented within the past 30 years. Lower Cretaceous strata of the Cedar Mountain Formation in eastern Utah have provided one of the most complete glimpses into this time period known in North America (Kirkland et al., 1998), whereas younger faunas are known mostly from southern Utah. For much of the Late Cretaceous, the Kaiparowits Basin located in the south-central part of the state, provides the most continuous and highest quality data in the region, and possibly the world (Eaton et al., 1999a). This trip is designed to give participants an overview of Cretaceous stratigraphy and paleontology of the Kaiparowits Basin, located primarily (Fig. 1) within Grand Staircase–Escalante National Monument, with an emphasis on recent discoveries and their implications. The Kaiparowits Basin area has been recognized for over two decades as containing one of the most continuous records of Cenomanian through Campanian terrestrial vertebrates in the world (Hutchison et al., 1997; Eaton and Cifelli, 1997). Although many workers have contributed to our understanding of Kaiparowits Basin paleontology, much of our present knowledge has resulted from two decades of reconnaissance level

sampling for microvertebrates via wet screenwashing of bulk rock samples by Rich Cifelli (Sam Noble Oklahoma Museum of Natural History) and Jeff Eaton (Weber State University). Faunal lists generated by Cifelli, Eaton, and others represent some of the highest documented Cretaceous vertebrate diversity for the Western Interior of North America (Eaton et al., 1999a; Foster et al., 2001; Sampson et al., 2004). Based largely on Eaton and Cifelli’s work (1997), the Late Cretaceous fossil record of southern Utah has come to be viewed as potentially key to resolving long-standing questions about continental-scale biogeography and diversity trends of everything from plants to dinosaurs (Archibald, 1997; Sampson et al., 2004). The importance of the Kaiparowits Basin is further underscored by its status as the only place known between Montana and the San Juan Basin of New Mexico to contain a diverse, well preserved, and continuous Late Cretaceous terrestrial biotic record. The stratigraphic sequence that hosts the fossil data is likewise impressive, comprising ~2000 m of clastic strata deposited in a broad range of environments in just over 25 m.y. In spite of the fact that the area was first mapped geologically over 125 years ago, many basic stratigraphic and paleontologic questions still remain unanswered. For vertebrate paleontologists, the region is still essentially a frontier, waiting to yield a wealth of data on the end Mesozoic biosphere.

Figure 1. Reference map showing field trip route, location of Grand Staircase–Escalante National Monument, regional Cretaceous outcrops, and political boundaries. Line of cross-section for Figure 3 shown as A–A′. Selected geologic structures also shown to delineate the Kaiparowits Basin.

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau REGIONAL OVERVIEW OF CRETACEOUS GEOLOGY The majority of Cretaceous outcrops (~850,000 acres) in the vicinity of Grand Staircase–Escalante National Monument are within the Kaiparowits Basin, a fan-shaped structural basin occupying most of eastern Kane and Garfield counties (Doelling and Davis, 1989). The Kaiparowits Basin is located at the eastern end of a 250-km-long outcrop belt of Cretaceous strata that stretches from the Pine Valley Mountains area near St. George to the Fifty Mile Cliffs in eastern Kane and Garfield counties (Fig. 1). The next outcrops to the east are found in the Henry Basin, separated from the Kaiparowits Basin by >40 km. The east limb of the Kaibab Anticline steepens sharply within a few kilometers east of the axis, forming the East Kaibab Monocline, which defines much of the western edge of the Kaiparowits Basin (Fig. 2). Steeply dipping Jurassic and Cretaceous formations in the East Kaibab Monocline have weathered into

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a spectacular series of hogback ridges known as “The Cockscomb.” Between the East Kaibab Monocline and Fifty Mile Mountain, Cretaceous units are mostly mildly deformed by a series of gentle north to northwest-trending folds, almost all of which are north plunging. The eastern limit of the Kaiparowits Basin is the axis of the Circle Cliffs Anticline, just west of the famous Waterpocket Fold. However, Cretaceous outcrops are found in the Kaiparowits Basin no farther east than the Sooner Bench, almost 75 km away. Two primary features distinguish the Kaiparowits Basin sequence. First, it contains the most complete upper Campanian-Maastrichtian sedimentary record in southern Utah (Fig. 3). Campanian rocks have not been unequivocally recognized west of the central Markagunt Plateau (see Nichols, 1997; and Eaton, 1999, for discussion), and upper Campanian rocks are thin or absent everywhere west of the Paunsaugunt Fault (Fig. 3). A relatively thick sequence of Campanian and Maastrichtian strata may

Figure 2. Generalized geologic map showing outcrops of Cretaceous units within the Kaiparowits Basin and surrounding area. Kdts—Dakota, Tropic, and Straight Cliffs formations; Kw—Wahweap Formation; Kk—Kaiparowits Formation; TKcg—Canaan Peak and Grandcastle Formations. Adapted from Doelling and Davis, 1989.

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A

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Kaiparowits Fm

Wahweap Fm.

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Masuk Mbr.-Beds on Tarantula Mesa* Muley Canyon Mbr.

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Canaan Peak Fm

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Morrison Formation San Rafael Group

Figure 3. Chronostratigraphic relationships of Cretaceous units in southern Utah. Note compression of stratigraphic units upsection because of regularly spaced time axis. No vertical thickness implied by units. Vertical ruling—unconformity.

also have been deposited in the Henry Basin, but if so, it, along with the Paleogene record, was later removed by post-Laramide (mostly Quaternary) erosion (Fig. 3). In southern Utah, Maastrichtian rocks are also largely confined to the Kaiparowits Basin. The second distinguishing feature of the Kaiparowits Basin Cretaceous sequence resulted from its paleogeographic position along the margin of the Cretaceous Western Interior Seaway for a 20 m.y. period. A complex interplay between eustasy, tectonics, and sedimentation caused several marine tongues to be interbedded with the largely terrestrial record, which has proven valuable for regional and global correlation of the terrestrial fossil record. In contrast, exposures in the Henry Basin record offshore marine conditions almost continuously from late Cenomanian to early Campanian time. A less complete Campanian nonmarine sequence is found in the Henry Basin, largely confined to the Tarantula Mesa area. West of the Kaiparowits Basin, the discontinuous stratigraphic distribution of fossils and lack of unequivocal marker beds in the primarily nonmarine section has left even the recognition of formational boundaries difficult, let alone precise chronostratigraphic correlation (Eaton, 1999; Eaton et al., 1999b).

to the Canadian polar region and east across the Great Plains. The Western Interior Basin developed in response to lithospheric flexure caused by thrust loading within the Sevier fold-andthrust belt coupled with eastward progradation of a thick clastic wedge. Because pre-Barremian Lower Cretaceous strata are rare or absent in Utah (Sprinkle et al., 1999), it is clear that the entire western margin of the southern Western Interior Basin was uplifted sometime between Tithonian and early Baremmian time. Some authors have attributed that uplift to the migration of a tectonic forebulge through the region (e.g., Currie, 1997; Willis, 1999). Many inconsistencies in age relationships between supposed pre- and post-forebulge deposits (particularly the absence of widespread pre-Barremian strata within both the eastern margin of the foreland basin and the supposed backbulge basin) make it unlikely that this model is entirely correct. It is fairly well established that in central Utah, thrusting during the Barremian resulted in deposition of the Cedar Mountain, San Pitch, and Indianola Formations (Sprinkle et al., 1999) in a distinctly foreland basin setting. However, there is no evidence that significant movement took place along the Sevier fault system in southern Utah or southern Nevada until Albian time (Carpenter, 1989).

TECTONIC–PALEOGEOGRAPHIC SETTING Late Early to Late Cretaceous (Albian-Maastrichtian) During the Cretaceous, the Kaiparowits Basin was located within the southern portion of the Western Interior Basin, a vast depositional trough that stretched unbroken from central Arizona

In latest Albian–early Cenomanian time, the classic frontal part of the Sevier Thrust Belt began to take shape as a continuous

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau

UPPER CRETACEOUS STRATIGRAPHY OF THE KAIPAROWITS BASIN The thick section of Cretaceous strata preserved in the Kaiparowits Basin (Fig. 5) is almost entirely mudstone, siltstone, and sandstone. Conglomerate, carbonate (siderite and calcite), coal, and phosphate are present in minor overall amounts. Five major source areas can be demonstrated by petrographic analysis. During initial sedimentation and periods of sediment bypass,

S G HL AN D HI C O RO G EN I SE VI ER

feature along the entire Wasatch Hingeline (DeCelles, 2004). A single crystal 40Ar/39Ar age of 97.9 ± 0.5 Ma (middle early Cenomanian, following Gradstein et al., 2004) was obtained for a sanidine collected from an ash immediately overlying the oldest Cretaceous foreland basin strata of the Zion area (Biek et al., 2003). This suggests that flexural loading of the region started no earlier than latest Albian. Coincident with thrust loading were the extremely high sea levels of the Greenhorn Eustatic Event (late Albian–middle Turonian). By late Cenomanian time, sandstone compositions in the Kaiparowits Basin indicate that granitic source areas in central Arizona were also making active contributions to the Kaiparowits Basin (Gustason, 1989). Although DeCelles (2004) shows a magmatic arc in southern Arizona during the Cenomanian, just exactly how the tectonics of this area affected the Kaiparowits Basin is still unclear. Elder and Kirkland (1993, 1994) postulate a rifted highland margin (Mogollon Highland) inboard of the Arizona magmatic arc as the source for Kaiparowits Basin feldspathic sands. Between Turonian and middle Campanian time, conditions remained similar to those established during the Cenomanian. Sea levels during the Niobrara Eustatic Event were elevated almost as high as those of the Greenhorn. Eastward movement of thrust plates continued. The influx of sediment into the area from the Mogollon Highland and Sevier Thrust Belt was balanced by the creation of accommodation space and high sea levels, allowing coastal conditions to persist in the Kaiparowits Basin region for almost 15 m.y. (Fig. 4). Thrust system fronts continued propagating eastward in middle and late Campanian time, and actually began to crosscut older Cretaceous foreland basin rocks of the Iron Springs Formation (Goldstrand, 1994; Lawton et al., 2003). Continued compression, coupled with an eastward sweep of magmatism in the late Campanian and into the Maastrichtian led to large-scale partitioning of the Western Interior Basin through regional uplifts like the San Rafael Swell and Circle Cliffs Dome, heralding the start of the Laramide Orogeny. Laramide phase activity along classic Sevier thrust fronts was accompanied by significant vertical offset on long-standing basement-seated structures like the East Kaibab Monocline and Waterpocket Fold. Locally thick, but laterally constrained basin fill sequences of Maastrichtian age characterize this phase of Kaiparowits Basin evolution. Continued regional uplift and the resulting regional unconformity between Cretaceous units and the upper portion of the Paleocene Grand Castle Formation draws the relevant part of the story to a close.

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INTERIOR SEAWAY

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Figure 4. Generalized paleogeographic map of the southwestern Western Interior Basin during the upper middle Turonian (lower Prionocyclus hyatti Ammonoid Biozone). Modified from Pollock (1999) and Elder and Kirkland (1993).

local sources were eroded and recycled. Most extrabasinal sediment can be traced to the Mogollon Highlands of central Arizona, the southern Utah portion of the Sevier Thrust Belt, the Delfonte Volcanic Province of southeastern California, or the southern Nevada portion of the Sevier Thrust Belt. Studies have demonstrated that even within a single formation there can be a complex interplay between all five sources, largely controlled by the contemporaneous regional tectonics (Eaton and Nations, 1991; Little, 1995; Lawton et al., 2003). Little (1995) divided the entire 2000 m column into four coarsening upward sequences (Fig. 5), relating them to periods of active tectonism (fine grained sequences) and post-tectonic quiescence (coarse gravel sheets and braided stream systems). Dakota Formation The lowest Cretaceous rocks in the Kaiparowits Basin are referred by most workers to the Dakota Formation (e.g., Eaton, 1991; Dyman et al., 2002; Ulicny, 1999). The Dakota is relatively thin in the Kaiparowits Basin (<60 m), especially compared to overlying formations. It consists of a heterogeneous mixture of mudstones, claystones, coal, sandstones, and chert pebble conglomerates that are divided into three informal members (Fig. 6): lower, middle, and upper (Eaton, 1991). The thickness of the member increases rapidly to the west, toward the axis of the foreland basin. East of Cedar City, it measures in excess of 300 m (Eaton et al., 2001). Environments are a mix of alluvial plain, coastal, and open marine as might be expected

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Per iod M Stag e ?U Su bsta CP Form ge atio n L Me mbe r

under the conditions of rapidly rising base level that characterized the Cenomanian (Kirschbaum and McCabe, 1992). Gustason (1989) and Titus (2002) demonstrated that syndepositional movement of crustal blocks within the Kaiparowits Basin strongly controlled lateral facies variability, thickness of

Absolute Ages

High-Res. Biostrat.

Significant Sequences Macrovert. (Little, 1995) Horizon

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74.21 Ma

2

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*

90.51 Ma

1 93.49 Ma

*

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Figure 5. Generalized stratigraphic column for the Cretaceous section exposed in the Kaiparowits Basin. Standard U.S. Geological Survey lithologic symbols used. Absolute dates with asterisks are based on analysis of bentonites associated with high resolution ammonite zonal indices and are from Obradovich (1993). J—Jurassic; V—various units; UJ—undifferentiated Jurassic; L—lower; M—middle; U—upper; TC—Tibbet Canyon; SH—Smoky Hollow; DT—Drip Tank; CS—Capping Sandstone; CP— Canaan Peak; M—Maastrichtian; Cen—Cenomanian; Con—Coniacian.

the formation, and diachroneity of the upper contact with the Tropic Shale in the middle and upper members. Lower Member Most workers refer the entire heterolithic Kaiparowits Basin sequence between the Jurassic-Cretaceous unconformity and the dark gray mudstones of the Tropic Shale to the Dakota Formation. Starting with Doelling and Davis (1989) it has been the practice of the Utah Geological Survey to refer the lowest portion of this sequence, which is largely conglomeratic, to the Cedar Mountain Formation. The primary lines of reasoning used to support this revision (Doelling and Davis, 1989; Doelling et al., 2003) are a supposed Early Cretaceous (Aptian-Albian) pollen date obtained from a carbonaceous interval in the middle of the lower member near Zion National Park and the presence of a gray-weathering smectitic layer resembling the upper portion of the Cedar Mountain Formation. The absence of angiosperms from the sample led Doelling and Davis (1989) to the conclusion that it was most likely pre-Albian or Aptian in age. However, Beik et al. (2003) report a date of 97.9 ± 0.5 Ma (middle early Cenomanian, following Gradstein et al., 2004) for an ash bed conformably overlying the conglomerate in the same general area. Because of the conformable nature of the contact between the conglomerate and overlying bentonite, we suspect the former is no older than late Albian. While the ash dates do show a time correlation of the conglomeratic unit with the Mussentuchit Member of the Cedar Mountain Formation, the Mussentuchit is not characterized by conglomerates. In contrast, upper Albian conglomerates (Nishnabotna Member) do characterize the basal portion of the Dakota Formation (Brenner et al., 2001) in its type area of eastern Nebraska and western Iowa. Furthermore, the middle member of the Dakota in the Kaiparowits Basin is quite smectitic, so the presence of a smectitic layer above the conglomerate in the Markagunt Plateau area is expected. Finally, the sub-Tropic units of the Kaiparowits Basin represent a fining upward retrogradational sequence resulting from increasing base levels (eustatic rise), not unlike the type Dakota. As a result, even though we acknowledge that the use of the term “Dakota Formation” in Utah may be inappropriate, we agree with Eaton (1991), am Ende (1991), Dyman et al. (2002), and Ulicny (1999) that the entire sub-Tropic Cretaceous interval of the Kaiparowits Basin should be referred to the Dakota Formation (Fig. 5). The lower member of the Dakota Formation is composed largely of white, gray, and black chert pebble, clast-supported conglomerate with subordinate chert lithic-rich sandstones, pebbly sandstones, and minor siltstones and mudstone lenses. Thickness for the unit in the Kaiparowits Basin ranges from 0 to 15 m. Thicker sections are located in the middle of paleovalleys. Massive bedding is common in the conglomerates, but trough and planar cross-beds are also regular features. Current indicators such as pebble imbrication and cross-beds demonstrate unimodal SE directed transport (Gustason, 1989), which is normal to the Sevier Thrust Belt. Upper Paleozoic fossils are common in the clasts and demonstrate derivation from the Sevier Thrust Belt.

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau Reworked pieces of older petrified wood are also common and are derived from the Chinle or Morrison Formations. Macrofossils other than reworked petrified wood are rare, although several relatively complete turtle shells were recently (summer 2005) found in gravelly facies. The age is constrained by the presence of ?Albian pollen and a middle early Cenomanian radiometric age date obtained from the base of the overlying middle member (see discussion above). Environmental indicators suggest deposition in braided stream systems of a high energy alluvial plane (Gustason, 1989; am Ende, 1991). Middle Member The middle member is a variable sequence of finer grained sandstone, massive and laminated mudstone, claystone, with occasional bentonitic clay layers, and carbonaceous beds. A wide variety of bedforms is present, including lenticular (isolated channel) bed sandstones, contorted bedding, multistory sandstone bodies, epsilon cross-beds, trough cross-beds, planar cross-beds, chevron cross-beds, ripple laminations, and horizontal laminations. Transport of sediment was largely east-southeast, away from the Sevier Thrust Belt (Gustason, 1989). The abundance of fines, as well as the isolated channels and other pronounced lateral accretion forms, indicate deposition in meandering and/or anastomosing stream systems and adjacent levees, swamps, and floodplains. Fossils are abundant locally, especially semionotid fish (Lepidotes) scales and teeth, crocodile teeth and scutes, turtle shell, bivalves, gastropods, and plant macrofossils, including foliage impressions and permineralized wood. All indicate freshwater, terrestrial habitats and a relatively humid climate. A bentonite sample taken from the middle portion of the middle member near the town of Tropic (Dyman et al., 2002) yielded a 40Ar/39Ar date of 95.97 ± 0.22 Ma, placing it firmly within the middle Cenomanian (following Gradstein et al., 2004), while the base of the member has yielded a middle early Cenomanian date over in the Zion area (Biek et al., 2003). Marine invertebrate biostratigraphy (see below) of the overlying member strongly suggests that the entire member is pre-late Cenomanian in age. Upper Member The upper member of the Dakota maintains some of the lithologic character of the lower Dakota, with sandstones, carbonaceous beds, and thick mudstone sequences. However, abundant feldspar grain content and the presence of persistent, thin bedded, hummocky cross-stratified sand sheets in the upper member allow it to be readily differentiated. Marine facies bioturbation and fossils of brackish water invertebrates such as Corbula, Brachiodontes, Crassostrea, Flemingostrea, and Exogyra are locally abundant. Less common are shells of open marine forms like Inoceramus and ammonites. Locally, the topmost bed contains a thick accumulation of massive-shelled oysters and other bivalves. A complex suite of backcoastal and coastal to offshore marine environments are indicated. Ammonites and inoceramids have demonstrated that the lower half is within the Calycoceras canitaurinum and Dunveganoceras problematicum biozones (Titus,

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2002). Fossils from the upper portion of the member refer it to the Metoicoceras mosbyense and, locally, Vascoceras diartianum ammonite biozones. This demonstrates near equivalency with the entire Hartland Shale Member of the Greenhorn Formation (lower upper Cenomanian). At least three transgressive-regressive parasequences can be recognized and traced over much of the Kaiparowits Basin (Ulicny, 1999). Tropic Shale The Tropic Shale (Fig. 7) was named by Gregory and Moore (1931). The type section is in exposures of dark gray marine shale found outcropping around the town of Tropic, Utah, just west of the Kaiparowits Basin. In the Kaiparowits Basin, the formation ranges in thickness from 183 to 274 m (Sargent and Hansen, 1982). Lithologies are predominantly mudstone and claystone, grading to more sandstone-rich toward the top of the formation. Solid and septarian carbonate concretionary horizons are characteristic of the lower and middle portions in the Kaiparowits Basin. However, concretions are also intermittently found in the upper portion. The upper and lower contacts are essentially conformable and both are defined as the change from mudstone, which characterizes the Tropic, to dominantly sandstone. Owing to the abundance of high resolution biostratigraphic data (ammonites and inoceramids) and co-occurring datable bentonite ash beds, the Tropic has been well constrained as late Cenomanian–early Turonian (Vascoceras diartianum-Prionocyclus hyatti ammonite biozones) and can be correlated virtually bed-by-bed with time equivalent marine strata in other parts of the Western Interior Basin. Elsewhere in Utah, Tropic equivalents have been referred to the Tununk Member of the Mancos, and the Allen Valley Shale of the central Wasatch (Hintze, 1988). The Tropic Shale was deposited in shallow water offshore to deeper water offshore muddy shelf facies. Titus (2004) recently reported the presence of cold hydrocarbon seep bioherms in the lower portion of the Tropic in the Cottonwood Canyon area. The mounds have up to 1 m of relief and occur over syndepositional upwarps that funneled methane from underlying coal beds in the Dakota Formation. Beautifully preserved fossils are abundant throughout the Tropic (Cobban, et al., 2000; Foster et al., 2001). Ammonites, inoceramid and mytiloidid bivalves, gastropods, and oysters are especially common, but corals, annelids (serpulids), and other taxa are known. A wide variety of small and large vertebrates, including sharks, bony fish, marine turtles, plesiosaurs, and a nearly complete dinosaur have also been collected (see description for Stop 2-1). Straight Cliffs Formation General The Straight Cliffs Formation (Fig. 5) was named by Gregory and Moore (1931—as the Straight Cliffs Sandstone), who described the type section as outcrops on the east face of the Kaiparowits Plateau (Straight Cliffs), south of the town of

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Figure 6. Photograph of the lower portion of the Cretaceous section of the Kaiparowits Basin. Shown are the Dakota Formation’s three members (lower, middle, upper) as well as the underlying Jurassic Entrada Formation and the overlying Tropic Shale and Straight Cliffs formations.

Figure 7. Photograph of exposures of Tropic Shale north of Big Water, Utah, in the vicinity of Stop 1-2. Important stratigraphic horizons in the lower portion of the Tropic are shown, including the Cenomanian-Turonian (C-T) boundary and the “B” and “C” bentonites of Elder (1991).

Escalante. The formation was further subdivided by Peterson (1969a) into, in ascending order, the Tibbet Canyon, Smoky Hollow, John Henry (Fig. 8), and Drip Tank Members. The type sections for all of the members are located in the south-central portion of the Kaiparowits Plateau, where exposures of each are especially good. The formation is mostly sandstone, with lesser amounts of interbedded siltstone, mudstone, and claystone. Coal and conglomerate are minor components. Because of the wide variety of coastal, offshore marine, and fluvial facies represented in the formation, the succession of sedimentary structures, bedding types, architecture, and lithologies is very complex. Fortunately, because the Straight Cliffs Formation contains the kind of large sandstone bodies that typically act as oil and gas reservoirs elsewhere, as well as rich coal resources, it has been the subject of several detailed studies (e.g., Doelling and Graham, 1972; Shanley and McCabe, 1993; Little, 1995; Hettinger, 1995; Castle et al., 2004). The formation ranges in thickness between 300 and 500 m, with a definite northward-thickening trend. Tibbet Canyon Member The Tibbet Canyon Member consists largely of tabular, crossbedded, and massive yellowish-tan sandstone and concretionary sandstone, with lesser amounts of thinly bedded gray and olive-green mudstone, particularly in the lower portion, where it conformably grades into the underlying Tropic Shale. Conglomerate is also a minor component. Conglomerates at the top of the member contain feldspar and nonrecycled monocrystalline quartz pebbles, indicating Mogollon source areas. Hummocky cross-bedding is common, as is planar and trough cross-bedding and marine facies bioturbation (e.g., Ophiomorpha). The lower contact is placed at the appearance of the first thick sandstone beds, but over much of the Kaiparowits Basin, it is essentially gradational with underlying Tropic Shale. The upper contact

Figure 8. Photograph of exposures of the upper portion of the Tropic Shale and lower three members of the Straight Cliffs Formation.

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau with the overlying Smoky Hollow Member is placed at the transition from predominantly sandstone to predominantly mudstone. Thicknesses range between 21 and 61 m (Cobban et al., 2000). Age constraints on the unit are entirely biostratigraphic, with ammonites and inoceramid bivalves of the Prionocyclus hyatti Ammonoid Biozone being fairly common in the lower half (Cobban et al., 2000). Marine to shoreface and backshore environments are indicated for much of the member in the Kaiparowits Basin, but the uppermost portion may record fluvio-deltaic conditions (Shanley and McCabe, 1995). The unit is highly fossiliferous, but vertebrate fossils are uncommon. Smoky Hollow Member The name Smoky Hollow Member was given by Peterson (1969a) to the carbonaceous mudstones, mudstones, coals, sandstones, and minor conglomerate that overlie the Tibbet Canyon Member and weather into a topographic bench. The designated type section (Peterson, 1969a) is in what is called Smoky Hollow on the U.S. Geological Survey (USGS) Nipple Butte 15 min quadrangle (1953 edition). When the area was remapped by the USGS at the 7.5 min level in the late 1970s, they changed place names, calling the former Smoky Hollow Squaw Canyon and naming an unnamed tributary of Warm Creek Canyon Smoky Hollow. Thus, the type locality is located in what is now called Squaw Canyon. The old road that used to access Squaw Canyon is now largely washed out, as well as administratively closed by Grand Staircase–Escalante National Monument, and the type section, as well as that of the John Henry Member, is difficult to access. Fortunately, nearby exposures in the present Smoky Hollow are equally good. In most places the unit is relatively thin (as little as 7 m) but it does thicken up to 100 m in the northeast part of the Kaiparowits Basin (Eaton, 1991). The lower contact with the underlying Tibbet Canyon is relatively sharp. The lower portion of the unit is relatively fine grained, frequently carbonaceous, while the middle portion, also mudstone dominated, tends not to be. An upper sequence consisting largely of coarse sandstone and coarse pebble conglomerate caps the member, and the contact with the overlying John Henry is also relatively sharp and marked by a transgressive surface. The upper sandy-conglomeratic interval was named the Calico Bed by Peterson (1969a), the name coming from its habit of weathering into a whitish or red and white mottled color that is very distinctive in outcrop. Locally, thick channel fill conglomerates that look identical to the Calico occur below the main mass of the Calico, separated from it by overbank mudstone intervals. The Smoky Hollow Member is not precisely age constrained largely due to a general lack of data for North American Turonian and Coniacian terrestrial vertebrate biostratigraphy. It is certainly not unfossiliferous; the middle noncarbonaceous interval yields abundant microvertebrate material. Regardless, the upper middle Turonian age of the underlying Tibbet Canyon Member and the middle Coniacian assignment given to the overlying John Henry Member constrains the unit fairly well (Eaton, 1991). A major sequence boundary (lower bounding surface of the Calico Bed)

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has been hypothesized by several workers (Hettinger, 1995; Shanley and McCabe, 1993; Castle et al., 2004). If this interpretation is correct, then the unit likely straddles the TuronianConiacian boundary. A rich microvertebrate fauna occurs in the middle noncarbonaceous interval (Eaton et al., 1999a) and the senior author has observed petrified wood (primary fossils, not reworked) and 2–3-cm-diameter bone fragments in the Calico Bed. A wide range of environments are interpreted for the member, including tidally influenced fluvial, estuarine, backswamp, and high energy braided fluvial. John Henry Member The John Henry Member is a heterolithic succession similar to the Smoky Hollow, but much thicker. Sandstone, carbonaceous and noncarbonaceous mudstone, and coal make up the bulk of the rock types. The type section, designated by Peterson (1969a), is just up Squaw Canyon, to the north, from the type Smoky Hollow, ~3 mi east of John Henry Canyon. As discussed under Smoky Hollow Member, Squaw Canyon was previously named Smoky Hollow on USGS maps. Thickness for the member ranges between 200 and 340 m, with a definite northeast thickening trend. Typical fluvial, shore-complex sedimentary structures characterize the John Henry including heterolithic epsilon-type and bidirectional crossbeds and hummocky cross-stratified sandstones. Coal is abundant in the member, particularly in what are known as the Alvey (upper) and Christensen (lower) zones. Seven separate mappable sand bodies, labeled A–G, have been recognized in the member (Peterson, 1969b). Fossils are locally abundant in the John Henry. As might be expected, faunas vary between fully terrestrial to fully marine, with brackish water forms such as sharks, rays, turtles (Eaton et al., 1999a), and oysters being especially common. The age of the John Henry is constrained by the occurrence of ?middle Coniacian ammonites and inoceramids in the lower portion and the diagnostic middle Santonian scaphitid ammonite (Clioscaphites) in an upper marine tongue (Cobban et al., 2000). Additional inoceramid data from the upper portion, from above the Clioscaphites horizon, suggest the uppermost portion is also upper Santonian (Cobban et al., 2000). Similar environments are interpreted for the John Henry as for the Smoky Hollow, including tidally influenced fluvial, estuarine, backswamp, and low and high energy fluvial. However, shallow offshore normal marine shelf deposits also occur in the unit. Drip Tank Member The Drip Tank Member consists mostly of resistant, tan to brown, massive, trough crossbedded and planar crossbedded sandstones, with locally abundant chert pebble conglomerate. The entire unit is characterized by amalgamated fluvial–multistory architecture. The member ranges in thickness from 43 to 150 m, with a N-NE thickening trend (Eaton, 1991). Because of its resistant nature, the Drip Tank locally weathers into vertical cliffs and underlies prominent benches. The lower contact with the John Henry appears lithologically gradational but is

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interpreted as an unconformity by most workers (e.g., Shanley and McCabe, 1993). Peterson (1969a) originally included finer grained units below a resistant sandstone ledge in his Drip Tank, but Eaton (1991) restricted the member to the resistant cliff-forming sandstone interval and placed the ledge forming upper portion of the type Drip Tank into the lower portion of the Wahweap Formation. The Drip Tank was deposited within high energy braided streams and possibly moderate energy meandering stream systems. Diagnostic macrofossils are not common in the Drip Tank, but Eaton (1991) has reported fish scales and fragments of turtle shell. This is not to say that vertebrate fossils are uncommon, as bone fragments up to 10 cm in length occur regularly in intraformational pebble lags above channel scours. Fossil wood, both silicified and limonitized, is abundant. Recent palynological sampling and analysis (Christensen, 2005) suggests the Drip Tank is entirely upper Santonian. Wahweap Formation General The Wahweap Formation (Fig. 5) was named by Gregory and Moore (1931—as the Wahweap Sandstone). The type locality is along the middle reach of Wahweap Creek, in the Kaiparowits Basin, where a thick section of resistant tan and yellowish-brown sandstones alternating with nonresistant olive-gray mudstones (Fig. 9) crop out between the Drip Tank Member and the overlying bluish-gray sandstones of the Kaiparowits Formation. In the Kaiparowits Basin, the formation ranges in thickness between 360 and 460 m. Eaton (1991) further divided the Wahweap Formation into four informal members, designating the Reynolds Point area, also in the Kaiparowits Basin, as the type section for the lower three, and Pardner Canyon as the type section for the uppermost because of accessibility. Contacts between the lower, middle, and upper members are all conformable, but sharp, and represent changes of sand/shale ratios and fluvial styles. The upper contact of the upper member with the overlying capping sandstone member represents an abrupt change in color, petrology, grain size and fluvial style. While there is some lateral variability in thickness, all four members can be consistently recognized throughout the Kaiparowits Basin. The lower three members are referred to the lower and middle Campanian (Aquilan Land Vertebrate Age) undifferentiated based almost entirely on the vertebrate fauna (Eaton, 1991). The capping sandstone member has been dated as middle Campanian based on palynomorphs (Pollock, 1999). The overlying Kaiparowits Formation has been dated at 76.1 Ma near its base giving a youngest possible age of middle Campanian. Overall the Wahweap Formation is highly fossiliferous, yielding a diverse lower and middle Campanian fauna. Lower Member The lower member is conformable with, but rests sharply on, the Drip Tank Member of the Straight Cliffs Formation. It forms alternating benches and slopes of interbedded quartzofeldspathic trough cross-bedded sandstone and mudstone deposited by N-NE–

flowing meandering rivers. In some areas of the southern Kaiparowits Plateau, sandstone becomes the dominant lithology, giving the member a less ledge-forming character. Eaton (1991) reported a thickness of 65 m for the unit at Reynolds Point, the type section. Sandstone composition and transport directions indicate Mogollon source areas were most significant. Intraformational conglomerates commonly overlie mudstone intervals above obvious scour surfaces. Petrified wood is common, and concentrations of large logs occur in several areas, particularly in southern outcrops. Eaton (1991) states that the highest concentrations of logs are on channel margins, associated with crevasse splay sandstones. Other fossils are also present, with stream-worn bone, turtle shell, and gar scales especially common in intraclast conglomerates concentrated as lags. Larger bone is not uncommon. One of the only true multitaxic dinosaur bonebeds known in the entire Kaiparowits Basin is in this member. The environment of deposition was in meandering stream systems and their adjacent floodplains on a middle alluvial plain (Eaton, 1991; Pollock, 1999). As discussed above, age constraints on the lower member are only general and are provided mostly by the Santonian age of the upper portion of the John Henry Member of the Straight Cliffs Formation. Middle Member The overlying middle member is a slope former, dominated by fine-grained deposits interbedded with quartzofeldspathic trough cross-bedded and ripple laminated sandstone. It is characterized by abundant point bar deposits and crevasse complexes deposited by meandering and anastomosing rivers that flowed to the north and northeast. Thickness is variable, but it generally measures ~100 m. Eaton (1991) reported a thickness of 112 m at Reynolds Point, the type section. Again, the high percentage of feldspar in the sandstones indicates significant Mogollon input. Intraformational conglomerate is common at the base of sandstone beds above scours. Mudstones are gray and green, showing occasional carbonate concretion development and other paleosol features. Fossils are common, and the member now has the distinction of containing the highest abundance of associated macrovertebrate fossil sites for the entire formation. Upper Member The upper member, which forms benches and ledges, is comprised of quartzolithic trough cross-bedded sandstone with minor mudstone deposited by northeast-flowing meandering rivers. Eaton (1991) reports a thickness of 148 m at Reynold’s Point. The member typically weathers into cliffs, making it difficult to prospect for fossils. Intraformational conglomerates above scours are common. The quartzolithic sandstone composition indicates a shift to sources in southern Nevada and possibly southwestern Utah. Eaton (1991) indicated the presence of large number of microvertebrate localities in the member. Capping Sandstone Member The capping sandstone member consists of multistoried conglomerate, pebbly sandstone, sandstone, and muddy siltstone

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sheets that represent braided gravel bars, channel complexes, crevasse splays, and floodplain deposits. Large siltstone and mudstone rip-up clasts (10–150 cm across) are commonly found at the bottoms of channel fill, which also contain abundant plant fragments and tree trunks with root networks. Thickness for the unit ranges from 75 to 150 m. Sandstone compositions are quartzose, and paleocurrents indicate deposition by E-SE–flowing braided rivers. The architecture, characterized by amalgamated channel deposits, has been interpreted as resulting from high energy braided stream systems (Pollock, 1999). The quartzose grain compositions of the sandstones strongly contrast with the feldspatholithic compositions of the lower two members, indicating a marked change in sediment source, presumably to the Sevier Thrust Belt. The unit is fossiliferous and the senior author has observed both large bone fragments and whole elements in the Henrieville Creek area. Kaiparowits Formation The Kaiparowits Formation (Fig. 5) was named by Gregory and Moore (1931). The type area was not designated, but its stated position above the brown sandstones of the Wahweap Formation and the lithologic descriptions given in a partial measured section make it fairly clear that the name was originally intended at least to include everything presently considered Kaiparowits Formation. Bowers’ (1972) designation of the Canaan Peak Formation clarified that the upper boundary of the formation was placed at a major regional angular unconformity, and the formation rapidly thins westward toward the Paunsaugunt Plateau. Where exposed under the Table Cliffs and Canaan Peak, the Kaiparowits Formation is the thickest Upper Cretaceous unit exposed in the Kaiparowits Basin. It is likely that the Kaiparowits Formation is not preserved anywhere west of the Paunsaugunt Fault (Lawton et al., 2003) even though it has been mapped there by previous workers. The formation consists mostly of gray-blue to greenish gray and gray sandstones, mudstones, and siltstones, and typically weathers into smooth-looking badlands (Fig. 10). Eaton (1991) reported a thickness of 855 m for the formation, which is almost identical to the 860 m recorded by Roberts et al. (2005). Recent work in the Kaiparowits Formation reveals the presence of three distinct units (lower, middle, and upper), which are characterized by changes in alluvial architecture, channel morphology, and sandstone/mudstone ratios. 40Ar/39Ar analysis of four bentonites collected from throughout the formation demonstrate a late Campanian age, between ca. 76.1 and 74.0 Ma (Roberts et al., 2005). The Kaiparowits Formation is abundantly fossiliferous with a diverse, entirely terrestrial fauna (DeCourten, 1978). Plant remains and freshwater invertebrate shells are found throughout. The fine-grained nature of the sediment and nonmarine fauna suggest deposition upon a low-relief, inland alluvial plain setting. Thick paludal deposits, large channels, and poorly developed, hydromorphic paleosols dominate the sedimentary record, and all are suggestive of a relatively wet alluvial system with periodic

Figure 9. The type sections for the Drip Tank Member of the Straight Cliffs (SC) Formation and the overlying lower, middle, and upper members of the Wahweap. View is to the northwest from the south side of Drip Tank Canyon, toward Reynolds Point.

Figure 10. “The Blues.” This is essentially the type section for the Kaiparowits Formation. View is looking north toward Powell Point (10,188 ft high) and the Table Cliffs Plateau. The topography is typical for outcrops of the Kaiparowits Formation.

aridity (Roberts, 2005). This is faunally supported by the high abundance and diversity of aquatic vertebrate and invertebrate fossils preserved within the formation. Canaan Peak Formation The Canaan Peak Formation (Fig. 5) is a relatively thin, coarse sequence of Sevier-derived chert pebble and cobble

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conglomerate and chert lithic–rich sandstones. A lesser amount of volcanic clasts are present. Trough cross-bedding is characteristic. High energy braided stream systems are the most characteristic depositional environment (Schmitt et al., 1991), and the size of the clasts indicate a not too distant source area for some of the sediment. Goldstrand (1994) suggested mixed source areas for the lower portion of the Canaan Peak Formation, with a component coming from southeastern Nevada (source of volcanic clasts) and a more local component derived from detritus shed eastward off of the Sevier Thrust Belt. The formation is not well age constrained and has been viewed by various authors as Maastrichtian, lower Paleocene, or straddling the K-T boundary. The Canaan Peak Formation has not yielded any diagnostic macrofossils and will not be observed on this trip. FIELD TRIP Day 1—Kanab, Utah, to Page, Arizona The first day’s trip will be an overview of the lower portion of the Cretaceous section. Two stops will be made. The first stop will focus on the nature of the Jurassic-Cretaceous unconformity and early Late Cretaceous (Albian-middle Cenomanian) sedimentation patterns and terrestrial vertebrate paleontology as recorded in the middle member of the Dakota Formation. The second stop will be an overview of the upper member of the Dakota Formation and the overlying Tropic Shale, which represent coastal and open marine environments. Directions to Stop 1-1 (from Kanab, Utah) Starting at the intersection of U.S. Hwy 89 with U.S. Hwy 89a (at the stoplight in Kanab), proceed east 46.6 mi (74.6 km) (Fig. 1). From there, turn left (north) onto the Cottonwood Canyon road. A monument information kiosk is located on the NW corner of this intersection. Proceed 0.9 mi (1.4 km) to the first cattle guard. Turn left immediately after the cattle guard into a loop pullout/turnaround two-track on the west side of the road. From the parking area, we will hike to the Jurassic-Cretaceous unconformity, located just south and east of the cattle guard, as well as over the small hill to the east to examine Dakota stratigraphy and a typical terrestrial vertebrate assemblage in outcrop. Please bear in mind that for most of the trip you are in a national monument and that all collecting of rock and fossil materials without a valid resource use permit is prohibited. Stop 1-2, which is not inside the Monument, will afford opportunities to collect marine invertebrates. West of the Paunsaugunt Fault, the highway is on the Upper Triassic Chinle Formation, while east of fault it is on the Lower Triassic Moenkopi Formation all the way to the East Kaibab Monocline. This whole portion of the drive is within the Grand Staircase physiographic province. As you pass east through the East Kaibab Monocline, the route cuts upsection through lower and middle Jurassic units. Some of the highlights along this drive include the monument boundary (mile 15.6), Paunsaugunt Fault (mile 15.7), the axis of the Kaibab Anticline

(mile 34.1), the Cockscomb–East Kaibab Monocline (mile 38.8), and Paria River (mile 43.1). Stop 1-1—The Rimrocks The Jurassic-Cretaceous unconformity is well exposed at this locality, where leached sandstones of the Middle Jurassic Entrada Formation are overlain by the middle member of the Dakota. Thin chert pebble lenses more characteristic of the lower member are found locally filling scours in the Entrada, especially to the south and east of the cattle guard. However, the lower member is missing over much of the immediate area. Only 12 mi to the east (19 km), it approaches 10 m in thickness within a wide NW-SE–trending paleovalley that can be traced all the way to the Butler Valley area. Locally, carbonaceous mudstone was deposited directly on the Entrada in basin shaped scours. Three multi-story channel fill sandstone lithosomes dominate the sequence near the road, separated from each other by relatively thin, fossiliferous olive-green to blue-gray overbank sequences and carbonaceous mudstones, probably of abandoned meander and paludal origin. Just a short hike over the hill shows that all three sandstones pinch out laterally into carbonaceous and noncarbonaceous mudstones. In this general area, the Dakota Formation is highly fossiliferous. Pollen and plant macrofossils indicate the flora was almost equally dominated by mosses and ferns, araucariacean and taxodiacean conifers (sometimes found as silicified logs), and a variety of angiosperms, including species of Magnoliaceae, Anacardiaceae, and Rosaceae (Gustason, 1989). Climate indicators such as fossil cycads (am Ende, 1991) and the abundance of unoxidized organic material suggest perpetually wet, warm, temperate to subtropical conditions. The abundant remains of tree stems and trunks preserved as limonitized external molds in fluvial channel deposits at this locality also underscore the lush nature of this ecosystem. Invertebrates, particularly freshwater ostracodes, gastropods, and bivalves are locally common. The recent collection of a well-preserved fossil aquatic insect nymph (cf. Odonata) clearly indicates that the potential this formation has to yield new information about middle Cenomanian biota is very high. Vertebrate body fossils and traces are found in a variety of Dakota lithofacies, with larger bones, scales, and teeth frequently concentrated above scours in simple or multistory sandstone bodies associated with intraformational conglomerates. Eaton et al. (1999a) reported 46 vertebrate taxa from the Dakota (mostly from the middle member) of the Kaiparowits area. This represents one of the most diverse middle Cenomanian faunas known anywhere in the world, yet Eaton et al. made it clear that study of the Kaiparowits Basin Dakota was still in its infancy. At this stop, crocodilians, turtles and amiid, and semionotid fish (Lepidotes) are the most obvious taxa. Their remains are found in at least three different lithofacies, including mudstones, intraclast pebble conglomerates, and chert pebble conglomerates. Most of the nonmammalian fauna remain unstudied by specialists, and even genus level assignments are still largely guesses. Eaton et

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau al. (1999a) listed four turtle taxa, the two most common of which are a form bearing granular ornament tentatively referred to the relict pleurosternid Glyptops, and a more lightly ornamented form tentatively referred to the baenid, Dinochelys. Glyptops is characteristic of Late Jurassic and Early Cretaceous faunas, and the Dakota occurrences represent the stratigraphically highest known. Although dominated by lacustrine and paludal fauna, the middle Dakota locally has produced primary type specimens for four early marsupials: Dakotadens morrowi Eaton, Alphadon clemensi Eaton, A. lillegraveni Eaton, and Pariadens kirklandi Cifelli and Eaton. In fact, Pariadens kirklandi held the record for the oldest known marsupial for several years (Cifelli and Eaton, 1987; Eaton, 1993). A diverse but poorly known dinosaur assemblage is also present, comprised of dromeaosaurids, ?troodontids, tyrannosaurids, nodosaurids, ankylosaurids, hadrosaurs, and hypsolophodontids (Eaton et al., 1999a; Kirkland and Parrish, 1995). This type of assemblage first appears in Utah in the upper portion of the Cedar Mountain Formation (Mussentuchit Member), in beds that are age equivalent to the lower member of the Dakota here in the Kaiparowits Basin. The overall aspect of the Dakota-Mussentuchit fauna is similar to those of Asia and highly dissimilar to sauropod-ankylosaur-iguanodontid–dominated faunas characteristic of the earlier Cretaceous of Utah (Eaton et al., 1997). Nokleberg et al. (2000) have shown that the formation of a continuous arc complex across the Bering Strait began in late Albian or early Cenomanian time, which seems to correlate closely with an influx of Asian exotics into North America and a continent-wide extinction of sauropods. Kirkland (1987) also documented some the youngest known occurrences of dipnoan lungfish in North America from the Dakota of this area. Tracks do occur in the middle member of the Dakota (Gustason, 1989), and recent field work by the Museum of Northern Arizona has documented a pentadactyl (probably ankylosaurid) dinosaur trackway in the roof of coal seam about 2 mi (3.2 km) west of this stop (Fig. 11). Directions to Stop 1-2 From the parking area for Stop 1-1, proceed south again, back across the cattle guard, to U.S. Hwy 89. Turn left onto 89 and drive east toward the town of Big Water. 10.4 mi (16.6 km) east of the junction of U.S. Hwy 89 with the Cottonwood Canyon road, turn left into the town of Big Water, by the Escalante Corner Mart. Proceed north 0.6 mi (1.0 km) until you merge with another paved road. Keep left, proceeding NW for another 1.3 mi (2.1 km), where the road curves sharply and descends through the middle Jurassic Camel Formation into the Wahweap Creek drainage. At the bottom of the hill, take the first right onto an obscure two-track. This intersection is at the south end of the fish hatchery. In 0.2 mi (0.3 km) after the turn, you will encounter a gate. Proceed through the gate and drive another 0.3 mi (0.5 km) to an abandoned gravel quarry. Park in the quarry and hike east across Wahweap Creek. Ascend the ridge of Dakota Formation before you (keeping south [right]) and then follow the ridge top to the north.

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Stop 1-2—Big Water Fish Hatchery This stop examines the upper marine–shoreline portion of the Dakota Formation and the overlying marine Tropic Shale. The contact between the two formations is especially well exposed in this area, as are some of the prominent marker beds within the Tropic, which require additional hiking to be observed up close. The upper member of the Dakota here is 17.5 m thick and is characterized by thin bedded sandstones and intervening mudstones, some of which are carbonaceous and/or coaly. Three distinct marine cycles can be recognized, each exhibiting a different succession of facies. In general, hummocky cross-stratified sandstones and gray or tan mudstones are characteristic of marine facies, while carbonaceous mudstones are characteristic of coastal mires. The lowest recognized flooding surface is characterized by a thin sandstone interval containing small Crassostrea and corbiculids and immediately overlies a thin carbonaceous interval. The top of the section is overlain by a 0.6-m-thick sandstone packed with the oyster Rhyncostreon levis (Stephenson). The Tropic Shale here is very well exposed for most of its entire thickness. Bentonites B–D of Elder (1991) are prominent, and limestones 3 and 4 (also of Elder, 1991) are represented by bands of concretions (Fig. 12). The lower half has yielded nearly all the macrovertebrate material, and all sites yielding diagnostic reptilian material are within the Mammites nodosoides or older ammonoid biozones. Several very significant macrovertebrate sites have been found in the lower half of the Tropic Shale in this general area (within a few miles) in the last eight years, many by local avocationals. Most of the sites have yielded plesiosaur remains, but marine turtles, fish, sharks, and a fairly complete dinosaur have also been found. No comprehensive summary of vertebrate diversity has been previously published for either the upper member of the Dakota or the Tropic Shale, although in the Kaiparowits Basin bones, teeth, and scales are commonly encountered. The upper Dakota has been observed to contain Squalicorax and other shark teeth and fish remains, but it is under-sampled. Recent fieldwork by the Museum of Northern Arizona and Northern Arizona University, supported by the Bureau of Land Management and National Park Service, has shown the Tropic contains a diverse fauna of chondrichthyans, osteichthyans, chelonians, and pliosaurid and polycotylid marine reptiles (Albright et al., 2002). Preservation is sometimes spectacular. Smaller bodied osteichthyan fish have been found perfectly articulated and maintaining three-dimensional body shape inside calcite concretions in the Cenomanian portion, although this is rare. A complete associated (probably originally articulated) tooth set identified as Ptychodus cf. decurrans Agassiz, and several well-preserved, associated or articulated specimens of large vertebrates, described below, have also been found. Among the more spectacular recent finds are the associated or articulated remains of large (3-m-long) xiphactinid fish (Fig. 13), at least four species of short-necked plesiosaurs, and three species of marine turtles. The plesiosaurs include three representatives of the family Polycotylidae: (1) the “common”

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late Cenomanian–early Turonian Trinacromerum; (2) a small, unnamed species that closely resembles a new taxon from South Dakota currently being described by other workers; and (3) a relatively large, probable new species of the genus Polycotylus, previously known only from the late Santonian–early Campanian. The family Pliosauridae is also represented by two large specimens of Brachauchenius lucasi Williston, both of which include skull material (Fig. 14) and one of which includes pectoral and pelvic elements previously unknown for this taxon. The marine turtles have not yet been studied in detail, but one appears to represent Desmatochelys lowi, whereas the other has a highly punctate-textured carapace that we assign to cf. Naomichelys. The most remarkable find made in the Tropic so far is the remains of a 70% complete therizinosaurid dinosaur (Gillette et al., 2001), tentatively referred to the genus Nothronycus. Return to U.S. Hwy 89, turn left, proceeding east to Glen Canyon Dam, and on to Page, Arizona, for the night. Day 2—Page, Arizona, to Escalante, Utah This leg of the trip will examine the coastal and alluvial plain deposits of the Straight Cliffs and Wahweap Formation exposed in the south-central portion of the Kaiparowits Plateau. Although the depositional history is highly variable, in general, this part of the section records a slow, episodic eastward withdrawal of the Cretaceous Western Interior Seaway and overprinted Sevier Tectonics. Three stops will look at paleontology and stratigraphy, while a fourth (Stop 2-3) will be made to look at the coal fire chimneys from which Smoky Mountain takes its name. Directions to Stop 2-1 From Page, drive west on U.S. Hwy 89 toward the town of Big Water. Turn right at the Escalante Corner Mart, proceeding north for 0.2 mi (0.3 km) to Utah State Hwy 12 (signed for Glen Canyon National Recreation Area). Turn right on Hwy 12, and proceed east for 12.6 mi (20.2 km), at which point you will see a fork in the road and a back country visitor kiosk. Take the right fork, and proceed east, dropping briefly into Warm Creek, and then ascending to the northeast, toward Smoky Mountain. At 13.6 mi (21.8 km), turn left onto the Smoky Mountain Road at the fork, maintaining a NE direction, heading toward Smoky Mountain. In a little under 3 mi you will begin the steep and narrow ascent of the infamous Kelly Grade. This road is not for the acrophobic! At mile 17.4, pull off to the right into a small alcove. Hike west into a small drainage, where you will be in the upper portion of the Tibbet Canyon Member. Be careful as the drainage leads to an exposed dry fall. Warning: Do not attempt to drive the Kelly Grade or hike this trail in wet or snowy weather. Stop 2-1—Lower Kelly Grade Excellent exposures of the upper portion of the Tibbet Canyon and Smoky Hollow members, as well as the lower portion of the John Henry Member, can be viewed at this stop (Fig. 15). Tibbet lithologies here are mostly medium and fine sandstone,

with thin mudstone laminae and beds also present. Trough and hummocky cross bedding is evident. Invertebrate traces are commonly found on bedding planes and cutting through beds. The entire sequence has been interpreted as nearshore in origin, with the amalgamated hummocky intervals generally interpreted here as lower shoreface and the trough and planar crossbedded sequences as upper shoreface. Near the type section of the Tibbet Canyon in Tibbet Canyon, a thin sequence of coarse pebble conglomerate and sandstone occurs at the very top of the member, suggesting possible fluvio-deltaic influence. The contact with the overlying Smoky Hollow Member here is typical, and sharp. Carbonaceous mudstones and coaly intervals rich with siderite are interpreted as coastal mires and are overlain by a coarsening upward sequence of middle and upper alluvial plain channel sandstones, conglomerates, and overbank fines. A lenticular lithosome of channel-fill pebble conglomerate almost identical in nature to the Calico Bed occurs in the middle of this sequence and demonstrates that high energy stream system facies pre-date the Calico Bed, the base of which is regarded by many to be a major sequence boundary. Shanley and McCabe (1991) use a eustatically driven sequence stratigraphic model to explain the depositional systems architecture observed in the Straight Cliffs, while others (Bobb, 1991; Eaton, 1991; Little, 1995; Lawton et al., 2003) infer episodic thrusting along the Sevier orogenic belt as the primary accommodation driver. In the latter model, high subsidence rates and accommodation in the proximal foreland basin during active thrusting tend to trap coarser sediment, allowing only mostly finer material to accumulate in the distal basin. During tectonic quiescence, rebound of coarser clastic sequences in the axial portion of the foreland basin permits the distal redistribution of coarser material. Undoubtedly, both eustatic and tectonic overprints are recorded in the sequence. The Tibbet Canyon Member is highly fossiliferous over much of its outcrop; however, the known fauna is mostly marine or brackish and is dominated by invertebrates. Eaton et al. (1999a) listed sharks, rays, gars, lungfish, crocodilians, and rare marsupial teeth from the member. The overlying Smoky Hollow Member is a different matter, and the senior author has observed bone fragments and silicified logs even in the high energy Calico Bed. Numerous microvertebrate localities, most of which are in the northern portion of the Kaiparowits Basin, have yielded a wealth of data. Eaton et al. (1999a) reported 61 taxa, with the mammals, dinosaurs, and squamates being especially diverse. Gardner (1999) named a new species of albanerpetontid amphibian, Albanerpeton cifellii, which was interpreted as being the most primitive member of its subgeneric clade known in North America. Nydam (1999) illustrated and described Dicothodon sp., a primitive teiid lizard from this member, and Voci and Nydam (2003) reviewed other parts of the squamate fauna. The only known pterosaur remains from the Kaiparowits Basin were also collected from the member (Eaton et al., 1999a). Cifelli (1990c), Eaton (1995), and J.G. Eaton (2005, personal commun.) described aspects of the mammalian fauna. The fauna has an entirely Late Cretaceous signature, with some of the Early

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau

Figure 11. Ankylosaur (cf. pes of Tetrapodosaurus) footprint mold from the roof of coal seam in the Dakota Formation, Rimrocks area of Grand Staircase–Escalante National Monument. The specimen is no longer accessible because of a roof collapse during the fall of 2004.

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Figure 12. Lower portion of the Tropic Shale at Stop 1-2. Shown are limestones 3 and 4 (LS-3, LS-4), the Cenomanian-Turonian stage boundary, and bentonites C and B of Elder (1991). Concretions of limestones 3 and 4 are highly fossiliferous. The interval around Bentonite “C” has produced several partial associated plesiosaur skeletons.

Cretaceous holdovers like Glyptops, Ceratodus, and Lepidotes conspicuously absent (Eaton et al., 1997). Directions to Stop 2-2 From the last stop, continue up the Kelly Grade for another 3.6 mi (5.8 km). Pull into a small pullout on your left, immediately after you cross a cattle guard. We will hike a short distance west and then drop down over the top to review the John Henry and Drip Tank Members of the Straight Cliffs Formation. This stop provides an incredible viewpoint looking south back toward Lake Powell and Page, Arizona. Stop 2-2—Upper Kelly Grade From this vantage you can see a section of the John Henry Member that is ~200 m thick. This general area was identified by McCabe and Shanley (1992) in their depositional facies models as transitional between raised coastal mires and fluvial and/or crevasse splay dominated depositional systems. Most of the member is interpreted as a highstand systems tract where raised mires may have regulated rates of change in accommodation space, creating an ideal environment for coal accumulation and stacked parasequences (McCabe and Shanley, 1992). The Kaiparowits Plateau has an estimated coal resource of 61 billion short tons (Hettinger, 2000). Most of the surface coal in this area has burned, leaving extensive stratigraphic breccias and red oxidized shale and sandstone behind. Eaton et al. (1999a) reported 42 vertebrate taxa from the John Henry Member. As might be expected, the formation yields abundant remains of marine and brackish water fish, and terrestrial components are rarer. A generic dromaeosaur-ankylosaur-hadrosaur dinosaur assemblage is known, and indicates the potential of the

Figure 13. Jaws of a large ichthyodectid fish (cf. Xiphactinus) collected from the lower Turonian portion of the Tropic Shale, Grand Staircase–Escalante National Monument. Occurrences of these kinds of fish are uncommon in pre-Coniacian rocks of North America.

member to yield important information on Coniacian-Santonian terrestrial vertebrates. No systematic attempts have been made to find vertebrate macrofossils in the member, although a number of ornithopod and theropod tracksites have been recently documented associated with coal seams (Foster et al., 2001). Here the Drip Tank Member consists of multistoried, pale yellowish brown to grayish orange, fine to coarse grained, well to poorly sorted, amalgamated channel sandstone that is trough and planar cross-stratified with interbedded quartz and chert pebble conglomerates and intermittent mudstones. Abundant log and stem casts occur above basal scours, along with minor occurrences

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Figure 14. Skull of the pliosaurid marine reptile Brachauchenius lucasi Williston collected in Glen Canyon National Recreation Area, east of Big Water, Utah.

Figure 15. Overview of the lower portion of the Straight Cliffs Formation seen at Stop 2-1.

Figure 16. Coal fire chimney on Smoky Mountain, Stop 2-3. Dark area in lower portion of photo is coated with a tar-like substance deposited around the smoking vent.

Figure 17. Leaf fossils (cf. Magnoliaceae) in the lower member of the Wahweap Formation collected from the Pilot Knoll area (Stop 2-4).

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau of dinosaur bone. Christensen (2005) recently examined the member’s composition and dispersal vectors at four locations within the Kaiparowits Basin. She identified detritus sourced from both the Sevier orogenic belt (quartzolithic) and Mogollon highlands (quartzofeldspatholithic). Christensen believes that N-NE–flowing longitudinal trunk systems (Sevier and Mogollon sourced) initially fed an underfilled basin during a time of active subsidence. During tectonic quiescence, the basin began to fill, transitioning to an eastsoutheast transverse directed dispersal system (Sevier sourced), and eventually creating a bypass surface and sequence boundary. This is contrary to work conducted by Shanley and McCabe (1995) who place a sequence boundary at the base of the unit and consider the Drip Tank a transgressive systems tract. No vertebrate faunal lists have ever been generated for the Drip Tank Member, and diagnostic material has not been studied. The quantities of bone seen in outcrops like these indicate the unit has future potential, although its tendency to weather into cliffs makes prospecting difficult. Directions to Stop 2-3 Drive 1.3 mi (2.1 km) north on the Smoky Mountain road. Turn right onto a spur road that heads east. Drive ~1.75 mi (2.8 km) on the obvious track, always staying to the right. One mile (1.6 km) in you will pass a drilling casing sticking about four feet out of the ground. Park and hike to the large smoking fissure, located south of the road. Stop 2-3—Big Smoky Coal Fire Naturally occurring coal fires have burned in this region for thousands of years, firing the landscape brick red. At this stop we will have a look at one of the three coal fires that continue to burn within the Kaiparowits Plateau (Fig. 16). The Big Smoky fire is located along the east limb of the Smoky Mountain anticline within the Alvey, Rees, and Christensen coal zones of the John Henry Member of the Straight Cliffs Formation. The fire was first reported back in 1951, yet a road was not constructed to this portion of the plateau until 1960. In 1967–1968 the U.S. Bureau of Mines (BOM) began efforts to try to extinguish the Big Smoky fire by drilling, blasting, and blading the area to smother active portions of fire. This effort was unsuccessful, and in 1976–1977, a similar approach was attempted, also without success. A study conducted during the 1998–2000 period suggested that the fire might be diminishing as it burns past the 400 ft depth. Directions to Stop 2-4 Return back from the coal fires via the route in. At the intersection with the Smoky Mountain Road, turn right and proceed north for 6.0 mi, until you reach the saddle of a small rock outcropping. Park wherever there is sufficient shoulder on the road. Stop 2-4—Pilot Knoll Pilot Knoll is a well known landmark along the Smoky Mountain road in the Ship Mountain Point quadrangle. This stop is meant to give an overview of the character, taphonomy, and

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faunal make up of the lower member of the Wahweap, which here has less mudstone than is typical elsewhere. Larger bone elements (mostly stream worn) are found at the bottoms of channel fill sandstones. Abundant fossil leaves are found inside of poorly cemented elliptical concretionary masses throughout the area (Fig. 17). Directions to Stop 2-5 From Pilot Knoll, drive north on the Smoky Mountain road. Be sure to keep right at the junction with the Head of the Creeks road (0.6 mi north of Pilot Knoll). The road skirts the east side of Ship Mountain Point and then descends into Drip Tank Canyon. The final descent into Drip Tank Canyon is steep and graded into ledge-forming sandstones that could potentially damage a lower clearance vehicle. Use caution when driving! At 3.9 mi north of Pilot Knoll, park to get an overview of the Wahweap Formation. Stop 2-5—Drip Tank Canyon Drip Tank Canyon affords a good opportunity to review the entire Wahweap Formation as it is the start of the type section for the lower three members (Fig. 9). The stratigraphic relationships, fluvial lithofacies associations, paleocurrents, and sandstone petrology of the Wahweap Formation (Fig. 18) here and elsewhere record alternating axial and transverse river systems in the southern Cordilleran foreland basin. The orientation of these river systems was controlled by the eastward advancing Sevier thrust belt during Late Cretaceous time. Although dating these nonmarine rocks is difficult and biostratigraphic data are limited, Pollock (1999) clarified the tectonostratigraphic relationships of the formation with an emphasis on the capping sandstone member. Fluvial rocks of the Wahweap Formation have been interpreted as tectonically dominated sequences that occupy the foredeep part of the foreland basin (Little, 1995). While that model fits the lower, middle, and upper members, the capping sandstone member represents a dramatic change in fluvial style, dispersal direction, and composition. This change marks an important transition in the history of the southern Cordilleran foreland basin Axial fluvial systems. Meandering and anastomosing rivers of the lower, middle, and upper members of the Wahweap Formation flowed parallel to the Sevier thrust belt during early Campanian time (Fig. 19A). These deposits are thick, overbankdominated successions deposited in a foredeep setting with high accommodation (Little, 1995). These rivers drained the Mogollon slope and highlands, an Early Cretaceous rift shoulder, and the southern part of the Sevier thrust belt that also contains thrusted volcanic rocks of a Jurassic volcanic arc. The lower members of the Wahweap Formation were deposited above alluvial and coastal plain deposits represented by the underlying Straight Cliffs Formation. Axial-fluvial systems occupy areas of maximum subsidence, proximal to the thrust front, parallel to orogenic belts (Miall, 1981). Transverse fluvial systems. The capping sandstone member represents a transverse-fluvial system flowing directly away from the thrust belt. The shift from axial to transverse

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Figure 18. Stratigraphic section for the Wahweap Formation at Reynolds Point, the type section for the lower, middle, and upper members. Quartz–feldspar–lithic (Qt-F-L) composition of sandstones given in percentages based on point counts. Arrows indicated mean fluvial transport direction.

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau

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Figure 19. Paleogeographic map of the southwestern U.S. during deposition of (A) axial river systems and (B) transverse river systems of the Wahweap Formation. NE-trending trunk river is hypothetical. NE-trending structures are fault-tip anticlines produced by eastward advancing thrust faults. KT—Kanarra thrust; ST—Sevier thrust; PT—Paunsaugunt thrust; EKT—East Kaibab thrust (from Pollock, 1999).

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dispersal was abrupt and is recorded by the provenance change and paleocurrent shift of the braided rivers that deposited the capping sandstone member of the Wahweap Formation during middle Campanian time (Fig. 19B). These sandy conglomeratic braided rivers are interpreted to record active uplift of Mesozoic sandstones above a frontal detachment and deformed in faultpropagation folds that provided quartzose detritus to the capping sandstone member. The abrupt change in dispersal direction and comparative thinness of the capping sandstone member combine to suggest that decrease in accommodation was combined with a change in paleoslope. These transverse braided rivers indicate that major uplift also took place in the interior thrust sheets of the Sevier thrust belt. Mixing of Precambrian-Cambrian quartzite clasts, Paleozoic cherts, and Mesozoic siliciclastic clasts derived from different thrust plates suggests that the conglomeratic part of the capping sandstone member was deposited by antecedent rivers that drained a wide region of the thrust belt. The single conglomeratic interval in the capping sandstone member represents separate uplifts of older, more interior thrust sheets, whereas the quartzose sandstones represent antecedent drainage that eroded Mesozoic eolian sandstones from fault-tip anticlines near the front of the thrust belt. The overlying Kaiparowits Formation represents yet another abrupt change in provenance, dispersal direction, and fluvial style and is restricted to the region around the Kaiparowits Basin. These changes have been suggested by Goldstrand (1991) to be indicators of the onset of Laramide deformation. Paleontology. The Wahweap Formation preserves the most diverse lower to middle Campanian terrestrial vertebrate fauna known in North America. Eaton et al. (1999a) and Cifelli (1990b) have documented four freshwater shark species, three freshwater ray species, seven bony fish species, two amphibian species, six turtle genera, two lizard taxa, three crocodilian taxa, eight dinosaur taxa, and 23 mammal species. However, the turtles, crocodilians, and dinosaurs require more complete skeletal material for specific identification. Over the past four years, the Utah Geological Survey, with funding from the Bureau of Land Management, has conducted a paleontological inventory of the lower sandstone and middle shale members of the Wahweap Formation in the southern Kaiparowits Basin (Kirkland, 2001). In addition to providing data on the distribution of paleontological resources, this study has identified and recovered specimens that add to our knowledge of larger vertebrates for this time interval, which is still poorly known. Although no significant crocodilian specimens have been found, fairly complete trionychid and baenid turtle shells have been recovered and are presently under study. Only two dinosaurs have been identified to species from rocks of this age in North America (both from Montana). At Grand Staircase– Escalante National Monument, cranial remains of a new species (Fig. 20) of long-horned centrosaurine ceratopsid have been the most significant dinosaur fossils to be identified so far. A number of associated hadrosaurid skeletons have been identified in the field, although taxonomically critical cranial remains have yet to

be identified in these as-yet preliminary excavations. One jugal does however compare well with Brachylophosaurus, a middle Campanian hadrosaurine from southern Alberta and Montana. The isolated skull roof of a juvenile pachycephalosaur (cf. Stegoceras) has been collected from the middle mudstone member. Additionally, carnivorous dinosaur remains have been identified at a number of sites, but nothing diagnostic has yet come to light. The most recent discovery is a problematic upper jaw that may represent a completely new group of animals not previously reported from the Wahweap Formation Day 3—Escalante to Henrieville Creek Culvert (The Blues) This day’s stops will examine the stratigraphy, depositional environments, taphonomy, and paleontology of the Kaiparowits Formation and depositional history of the Capping Sandstone Member of the Wahweap Formation. The thick succession preserved in the large amphitheater of badlands known as “The Blues” record a diverse and well preserved fauna and shifting sediment source areas as the Sevier phase of the Cordilleran Orogeny drew to a close. Directions to Stop 3-1 From the monument visitor center turnoff, located on the west edge of town, drive west on Utah State Hwy 12 toward Henrieville for 17.8 mi (28.5 km) to a developed overlook/pullout. There is a toilet and interpretive sign, as well as a small paved parking area. Stop 3-1—The Blues–Powell Point Overlook Stratigraphy. Recent facies and architectural analysis of the Kaiparowits Formation has begun to flesh out details of the formation’s paleoenvironments and depositional history (Roberts, 2005). The Kaiparowits Formation is subdivided (following Roberts, 2005) into three informal units: upper, middle and lower, based on distinct changes in alluvial architecture, channel morphology, and sandstone/mudstone ratios (Fig. 21). The lower and upper units are characterized by high channel/overbank ratios of 75/25 and 60/40, respectively, sheet-like major sandstones, and only moderately abundant fossil preservation. In contrast, the middle unit is richly fossiliferous, contains a high proportion of lenticular channels, and a low channel/overbank ratio (45/55). Paleoenvironmental analysis demonstrates that thick paludal deposits, large channels, and poorly developed, hydromorphic paleosols dominate the sedimentary record. All are suggestive of a relatively wet alluvial system with periodic aridity (Roberts, 2005). Such assertions are faunally supported by an elevated abundance and diversity of aquatic vertebrate and invertebrate fossils preserved within the formation. New laser-fusion 40Ar/39Ar analysis of four bentonite horizons has also resulted in the first absolute ages for the 860-m-thick Kaiparowits Formation (Roberts et al., 2005), resolving previous age uncertainty (see Eaton, 1991). A late Campanian (Judithian) age, between ca. 76.1–74.0 Ma, is assigned to the formation.

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau

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Figure 20. Cross section of a long horned ceratopsid skull (cf. Centrosaurinae) recently collected from the Wahweap Formation near the transition of the lower sandstone to the middle mudstone member, Grand Staircase–Escalante National Monument. Visible are the occipital condyle, orbital horns (top center), frill (left center), maxilla, and facial bones.

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These new ages reveal an extraordinarily high sediment accumulation rate (41 cm/k.y.) for the Kaiparowits Formation, which ranks among the highest reported for any fluvial and floodplain sequence in the Western Interior Basin. The unusually thick and rapidly deposited nature of the formation are interpreted to be the result of a wet paleoenvironment and sub-humid climate, coupled with rapid subsidence and generation of accommodation space during syntectonic thrust loading, local (tectonically driven) base level rise, and possibly even the onset of Laramide uplifts within the foreland. Detailed stratigraphic correlation reveals that the Kaiparowits Formation is contemporaneous with many of the most important vertebrate fossil–bearing formations in the Western Interior Basin (see Roberts et al., 2005), which provides a new chronological basis for addressing questions relating to mammal biostratigraphy and vertebrate evolution, biodiversity, and paleobiogeography (e.g., dinosaur provincialism) in the Cretaceous Western Interior Basin. In addition, the Kaiparowits Formation can now be more accurately correlated with other well-studied strata across Utah and southeastern Wyoming, including portions of the Book Cliffs sequence. Paleontology. The Kaiparowits Formation is considered by many to be the “crown jewel” of Kaiparowits Basin paleontology because of the abundance, diversity, and quality of its fossils. Previous research has conclusively demonstrated the formation contains an outstanding record of Late Campanian terrestrial vertebrate ecosystems. Eaton et al. (1999a) reported 88 vertebrate taxa, this list being derived largely from screenwashing of mudstones for micro material (see also Cifelli, 1990a, 1990c,

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Figure 21. Composite measured section of the Kaiparowits Formation in Grand Staircase–Escalante National Monument (from Roberts, 2005). Location and 40Ar/39Ar ages of the four bentonites (ash beds) are shown.

1990d; Eaton, 1991, 2002; Nydam, 1997; McCord, 1997, 1998; Gardner, 1999). With regard to macrofossils, several institutions, including the Museum of Northern Arizona, Brigham Young University, and University of California–Berkeley, have conducted sporadic work (DeCourten and Russell, 1985; Weishampel and Jensen, 1979; Parrish and Eaton, 1991; Hutchison, 1993). These pioneering efforts have demonstrated that the formation also has the potential to yield diverse vertebrate macrofossils. Since much of the previously studied material was derived from screenwashing, it is difficult to diagnose to genus (with the exception of the

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mammals and squamates). This also makes it difficult to make meaningful comparisons with other faunas. Recent fieldwork by the University of Utah, the Raymond Alf Museum (Claremont, California), and the Museum of Northern Arizona has specifically targeted macrofossils to increase or refine known vertebrate diversity of the Kaiparowits Formation. Below is a taxon-based summary of the paleontology of this highly productive unit, including summaries of recent finds, and a brief review of some biogeographic implications. Fishes. A total of 13 genera of fish have been identified from the Kaiparowits Formation (Eaton et al., 1999a). The chondrichthyan fauna includes the primitive hybondont genus Hybodus and the shell crushing Lissodus, known from nearshore marine and freshwater habitats respectively; orectolobiform carpet sharks Brachaelurus and Squatirhina; the sclerorhynchid sawfish Ischyrhiza; and the rhinobatid guitarfish Myledaphus. Bony fish are also well represented and include the gar Lepisosteus, the bowfins Amia and Melvius, and the sturgeon Acipenser. More advanced teleosts include Paralbula, Paleolabrus, and Plactacodon. Amphibians. Amphibians make up only a minor portion of the known microvertebrate fauna in the Kaiparowits Formation; however, a total of eight genera have been recovered, including representatives from both the Lissamphibia and Anura. The two most abundant genera are Albanerpeton and Habrosaurus (Eaton et al., 1999a; Gardner, 1999, 2000; Nydam, 2004, personal commun.), which are interesting biogeographically because Albanerpeton is found only in the Kaiparowits Formation and in more southern formations within the Western Interior Basin (Armstrong-Ziegler, 1980). Conversely, Habrosaurus is known from the Kaiparowits Formation and more northern formations within the Western Interior Basin. Squamates. Recent work on squamates has raised the generic diversity to 13 taxa (McCord, 1997; Eaton et al., 1999a; R.L. Nydam, 2005, personal commun.; Voci and Nydam, 2003). Teiid lizards are the most common squamate fossils, but scincids, xenosaurids, helodermatids, anguids, veranids, and a species of snake are also represented (McCord, 1997, 1998; Eaton et al., 1999a; R.L. Nydam, 2004, personal commun.). Turtles. Turtles fossils are relatively common throughout the Kaiparowits Formation. The most common taxa are the genera Basilemys, Adocus, “Baena,” Aspideretes, and Compsemys (Eaton et al., 1999a; Hutchison et al., 1997). Trionychids are the most common turtles found in the Kaiparowits Formation. Other components of the turtle fauna include a small, undescribed mud turtle and Cretaceous members of the Chelydridae, or snapping turtles. Crocodilians. Crocodilian remains are both abundant and diverse in the Kaiparowits Formation; yet this group remains poorly known due to the fragmentary nature of most specimens (Eaton et al., 1999a). The formation possesses the most diverse crocodilian fauna known in the entire Western Interior Basin, with five taxa identified, including an undescribed goniopholid, an undescribed species of Brachychampsa, and a new genus of caiman, the oldest occurrence of this group discovered to date (Wiersma et al., 2004). Two of five genera were found in the last

four years. The diverse morphologies of the coeval Kaiparowits taxa suggest equally diverse lifestyles. Dinosaurs. Hadrosaurs are by far the most commonly recovered dinosaur in the Kaiparowits Formation. However, until 2004, only one taxon had been identified even to genus level, the crested lambeosaurine Parasaurolophus cyrtocristatus, otherwise known only from New Mexico (Ostrom, 1961; Weishampel and Jensen, 1979; Sullivan and Williamson, 1999; Gates, 2004). In 2004, two exquisitely preserved, articulated hadrosaurine skulls were recovered. Both specimens are tentatively assigned to the same taxon, a new species of Gryposaurus (Fig. 22). Many other elements of partial skulls have also been collected, some associated with relatively complete post-cranial skeletons. Several of the articulated skeletons found exhibit internal and external molds of soft parts including skin impressions (Fig. 23). Hypsilphodontids are also well known and a specimen tentatively assigned to this family is the best-preserved dinosaur yet recovered from the Kaiparowits Basin. Pachycephalosaur remains in the Kaiparowits Basin, like elsewhere, are rare. However, several isolated specimens have been recently recovered, including two highly ornamented squamosals from a small-bodied taxon similar to Stegoceras. Ceratopsids are also known, and three recently discovered localities have yielded remains of a new genus of chasmosaur, which represents the first new genus from this subfamily recognized in more than 75 years (Smith et al., 2004). One specimen consists of a mostly complete skull (Fig. 24) and postcranium (Getty et al., 2003). Phylogenetic analysis (Smith et al., 2004) demonstrates that this new animal forms a clade with Pentaceratops and “Chasmosaurus” mariscalensis, which in turn is the sister group of a second clade composed of the three northern species of Chasmosaurus (all from Dinosaur Provincial Park in Alberta). In contrast to earlier studies (e.g., Forster et al., 1993), these results provide strong support for the dinosaur provincialism hypothesis (Lehman 1987, 1997, 2001) and suggest the possibility of separate centers of endemism within the Western Interior Basin during the Campanian (Sampson et al., 2004). Ankylosaurs comprise a small component of the Kaiparowits Formation dinosaur fauna. However, the Kaiparowits Basin record includes several associated specimens recovered in the last four years. One locality has yielded >40 associated osteoderms of varying sizes from a nodosaurid (cf. Edmontonia), while Eaton et al. (1999a) report the presence of the ankylosaurid Euoplocephalus. Other ankylosaur material has also recently been found. Small and large-bodied theropods are also known. All of the larger materials found to date can be assigned to Tyrannosauridae, in keeping with other Campanian deposits in the Western Interior Basin (Carr and Williamson, 2000, 2005). Finds of isolated tyrannosaur teeth are quite common, and several localities have produced partial skeletons. An associated tyrannosaurid specimen collected in the 1980s by a crew from Brigham Young University, and now under study by Thomas Carr (Carthage College), represents a new genus and species (T. Carr, 2005,

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau personal commun.). Small theropod diversity includes dromaeosaurids, troodontids, ornithomimids, oviraptorosaurs, and the bird Avisaurus. The first record of oviraptorosaurs consists of a nearly complete articulated hand and partial foot of a caenagnathid, representing a new genus and species and a dramatic southern range extension for the group in North America (Zanno and Sampson, 2003). Mammals. Mammals are one of the best understood vertebrate groups from the Kaiparowits Formation. For example, at least 13 mammal genera have been identified from this unit, including representatives from multituberculates, marsupials, and insectivores (Cifelli, 1990a, 1990b, 1990d; Eaton, 1995; Eaton et al., 1999a). Of these, four genera—two multituberculates (Cedaromys and Kaiparomys), one marsupial (Aenigmadelphys), and one placental (Avitotherium)—are endemic to the Kaiparowits Formation. Interestingly, despite two decades of intensive field work, pediomyid and stagodontid mammals, both of which occur in the underlying Wahweap and Straight Cliffs Formations, are absent from the Kaiparowits Formation (Cifelli, 1990a, 1990b; Eaton et al., 1999a). Biogeographic implications. The Kaiparowits fauna is essentially time equivalent (Roberts et al., 2005) to those found in other dinosaur-rich formations to the north (Dinosaur Park Formation, Alberta; upper portions of the Judith River and Two Medicine Formations, Montana), east (Fruitland Formation and the lower portion of the Kirtland Formation, New Mexico), and south (upper shale member of the Aguja Formation, Texas). Because of the unusual amount of endemism observed in the Kaiparowits, Sampson et al. (2004) hypothesized that, despite mild climates and small landmass size, terrestrial faunas of the Late Cretaceous Western Interior are highly provincial, perhaps even more than previously observed. Regional endemism in terrestrial vertebrates seems to parallel that described for Campanian marine invertebrate and vertebrate faunas from the adjacent Cretaceous Western Interior Seaway (Kauffman, 1984; Nicholls and Russell, 1990), leading to the possibility that controls were mostly latitudinal temperature variation (Kauffman, 1984; Nicholls and Russell, 1990; Lehman 1987, 1997, 2001), especially since evidence for major topographic barriers is lacking. Exactly how this translates onto land is not clear, but pollen data clearly demonstrate there are two distinct floral realms. If disjunct vegetative communities persisted long enough, northern and southern herbivores, and their predators, may have specialized and diverged accordingly, producing the apparent biogeographic pattern. Directions to Stop 3-2 From the overlook, drive west on State Hwy 12 down the dugway ~0.3 mi (0.5 km). Park on the shoulder and take a small walk to examine newly discovered insect nest fossils. Stop 3-2—The Blues–Upper Grade This is the only known locality for a new continental trace fossil (Fig. 25), and so far nine discrete nest structures have been identified (Roberts and Tapanila, 2005). They are all preserved in

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a single bedding plane, within an area of ~25 m2. Fine-grained, ripple-laminated sandstone, interpreted as a fluvial crevasse splay deposit, drapes the nest structures and is thought to have rapidly buried and preserved much of the subaerial morphology of the nest structures. Ants and termites are considered to be the most likely tracemakers, based on the complex network of unlined, cylindrical galleries and chambers that are located within the conical mound structure. Such dense arrangements of branching and anastomosing galleries and more rare chambers are the hallmark of modern social insect nests (e.g., Hasiotis, 2003). This trace fossil preserves rare, tangible evidence of nest construction by social insects during the Late Cretaceous and provides an important new datum for reconstructing the evolutionary and ecological history of social insects. Since nests commonly have multiple functions, including shelter, food and garbage storage, and egg and young rearing, they can provide rare insight into the behavior and social structure of the tracemaker. Such fossil insect nests as the new Kaiparowits ichnotaxon provide direct evidence tracing back the antiquity of social behavior in the fossil record. Analysis of these complex nest structures also aids in paleoenvironmental reconstruction, revealing multiple phases of nest construction, burial, and reestablishment within an active floodplain system (Roberts and Tapanila, 2005). Directions to Stop 3-3 Drive an additional 0.8 mi westbound on Hwy 12, and pull off onto a small pullout on the east (left) side. Hike a short distance north, back into the middle unit of the Kaiparowits Formation, to examine two of the most common types of vertebrate fossil taphonomies in the formation: isolated large elements (unassociated macrosites) and channel fill small bone lags (fluvial microsites). Stop 3-3—The Blues–Middle Grade Based on an analysis of 276 vertebrate fossil localities, four primary modes of fossil preservation are recognized in the Kaiparowits Formation: microsites, isolated and/or unassociated macrosites, associated macrosites, and articulated macrosites (Roberts et al., 2003; Roberts, 2005). Microsites and isolated and/or unassociated macrosites represent the most common mode of preservation in the formation and are useful for evaluating broad trends in taxonomic diversity and mammalian biostratigraphy. Articulated and associated macrosites are much rarer in the formation, but are important for studies dealing with the detailed taxonomy and ecological and evolutionary relationships of the fauna. In addition, an unusual variety and abundance of taphonomic features is observed on bones from throughout the formation, including “wet rot” weathering, carnivore tooth traces, trample traces, and insect traces (Roberts et al., 2003). The relationships between taphonomy and lithofacies demonstrates a complex and diverse Late Cretaceous ecosystem for the Kaiparowits Formation and suggests that trends in preservation were strongly influenced by a relatively wet paleoenvironment and a humid to sub-humid

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climate that was subject to periodic (perhaps seasonal?) aridity. Preservational style and quality of fossils varies significantly between channel and overbank facies. Overbank facies are dominated by poorly preserved, isolated and associated macrosites with copious bone modification features, such as “wet-rot” bone weathering, carnivore tooth traces, trample traces, and insect

borings, which are likely associated with rapid subaerial decay and decomposition on the humid floodplain. In contrast, channel facies preserve a range of poorly preserved, isolated macrosites and microsites to exquisitely preserved, fully articulated skeletons (single individuals) with soft tissue impressions, indicative of both attritional and rapid burial, but not mass mortality. The paucity of bone beds in the formation may be attributed to generally equitable climatic conditions, where typical agents of mass mortality, such as drought, famine, and flash-flooding, were rare. Directions to Stop 3-4 From the last stop, drive an additional 1 mi and pull off to the south (left) on a broad pullout just before the road crosses a large drainage culvert. Hike west from the pullout into the

Figure 22. Excellently preserved articulated skull of the hadrosaurin dinosaur Gryposaurus collected summer 2004 from the Kaiparowits Formation, Grand Staircase–Escalante National Monument.

Figure 23. Well preserved internal molds of epidermis from the dorsal area of the mid-caudal region of a hadrosaurine dinosaur. Specimen was collected summer/fall 2004 from the Kaiparowits Formation, Grand Staircase–Escalante National Monument. Impressions were preserved in life position, closely associated with articulated skeletal fossils.

Figure 24. Reconstructed skull of a new genus and species of chasmosaurine ceratopsid collected in 2003 from the Kaiparowits Formation, Grand Staircase–Escalante National Monument.

Figure 25. New social insect nest ichnotaxon from the Kaiparowits Formation. Shown is a bedding plane view of the subaerial portion of the nest structure, which is draped by a ripple-laminated crevasse splay deposit.

Stratigraphy, environments, and paleontology of the Kaiparowits Plateau bottom of the creek and follow it under the road (north) into an arroyo cut into Kaiparowits Formation to examine depositional modes and processes. Stop 3-4—Henrieville Creek Culvert Based on recurrent patterns of grain size and lithology, internal and external geometry, sedimentary and biogenic structures, and paleontology, strata in the Kaiparowits Formation have been subdivided into nine distinct, repeated lithofacies assemblages (Roberts, 2005). The focus of this stop is to observe one of the more unique lithofacies assemblages in the formation: mollusc shell conglomerates. Mollusc shell conglomerates are easily identifiable but rare in the formation. They are composed of densely packed, fluvial bivalve shells, referable to Unio sp. and Plesielliptio sp. (see McCord, 1997), in a medium-grained sandstone matrix (Roberts, 2005). The concentration of unionid shells varies vertically and horizontally within individual units. Thickness of beds ranges from 0.3 to 2 m, and individual beds extend laterally between 7 and 50 m, commonly grading into other lithofacies assemblages, such as intraformational conglomerates or major sandstones. At least three mollusc shell conglomerates exceeding 1 m, and multiple thinner beds, have been identified in the Kaiparowits Formation. In the bed examined at this stop, three distinct intervals are observed: (1) A basal sandy unit above a fourth-order bounding surface, composed of shell-hash in a coarse sand and mud pebble matrix; (2) a distinctive middle interval of densely packed, commonly articulated unionid shells, with minimal sand matrix; and (3) a gradational upper unit composed of shell fragments and few articulated shells in a medium sand matrix. Unionid clams are known to live in high concentrations in shallow, sandy fluvial environments, particularly within shallow pointbar deposits; however, densely packed mollusc shell beds are uncommon within the nonmarine record, particularly in fluvial settings. This and other similar unionid shell beds recorded in the formation likely represent rare mass mortality concentrations, as indicated by the majority of closed, articulated shells. Similar bivalve census deposits have been observed along marine shorelines immediately following tropical and winter storm events (e.g., Boyajian and Thayer, 1995). Intense winnowing of finer sand during storms results in dense shell concentrations. The subsequent inability of the bivalves to dig back into the substrate leads to mass mortality and death assemblages. Since the architecture of the Kaiparowits shell beds is virtually identical to shell beds described by Boyajian and Thayer (1995), a similar genetic relationship is implied. ACKNOWLEDGMENTS All of the authors have been either directly or indirectly supported by Grand Staircase–Escalante National Monument and the Bureau of Land Management. In particular, Marietta Eaton and Dave Hunsaker, both of Grand Staircase–Escalante National Monument, were key to the success of the last five

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years. Randy Nydam, J. Howard Hutchison, Don Lofgren, Jeff Eaton, and Amy Christensen have provided critical discussions and/or faunal data. Numerous volunteers, including many from the Utah Friends of Paleontology, have assisted with aspects of field and laboratory work. REFERENCES CITED Albright, L.B., III, Gillette, D.D., and Titus, A.L., 2002, New records of vertebrates from the Late Cretaceous Tropic Shale of southern Utah: Geological Society of America Abstracts with Programs, v. 34, no. 4, p. 5. am Ende, B.A., 1991, Depositional environments, palynology, and age of the Dakota Formation, south-central Utah, in Nations, J.D., and Eaton, J.G., eds., Stratigraphy, depositional environments, and sedimentary tectonics of the western margin, Cretaceous Western Interior Seaway: Geological Society of America Special Paper 260, p. 65–83. Archibald, J.D., 1997, Emerging importance of the Grand Staircase–Escalante region in Cretaceous vertebrate biostratigraphy, western U.S., in Hill, L.M., ed., Learning from the land—Grand Staircase–Escalante National Monument Science Symposium Proceedings: U.S. Department of the Interior, Bureau of Land Management, p. 355–357. Armstrong-Ziegler, J.G., 1980, Amphibia and Reptilia from the Campanian of New Mexico: Fieldiana Geology, p. 1–39. Biek, R.F., Willis, G.C., Hylland, M.D., and Doelling, H.H., 2003, Geology of Zion National Park, in Sprinkel, D.A., Chidsey, T.C., Jr., and Anderson, P.B., eds., Geology of Utah’s Parks and Monuments: Utah Geological Association Publication 28 (2nd edition), p. 107–137. Bobb, M.C., 1991, The Calico Bed, Upper Cretaceous, southern Utah; a fluvial sheet deposit in the Western Interior foreland basin and its relationship to eustasy and tectonics [M.S. thesis]: Boulder, University of Colorado, 166 p. Bowers, W.E., 1972, The Canaan Peak, Pine Hollow, and Wasatch formations in the Table Cliff region, Garfield County, Utah: U.S. Geological Survey Bulletin 1331-B, 39 p. Boyajian, G.E., and Thayer, C.W., 1995, Clam calamity; a recent supratidal storm deposit as an analog for fossil shell beds: Palaios, v. 10, p. 484–489. Brenner, R.L., Ludvigson, G.A., Witzke, B.J., Joeckel, R.M., Phillips, P.L., Gonzalez, L.A., and Ufnar, D.F., 2001, Deposition of fluvial gravels in late Albian Kiowa marine cycle fluvial-estuarine setting, Nishnabotna Member, Dakota Formation along the cratonic margin, Cretaceous Western Interior Basin: Geological Society of America Abstracts with Programs, v. 33, no. 6, p. 355. Carpenter, D.G., 1989, Geology of the North Muddy Mountains, Clark County, Nevada, and regional structural synthesis; fold-thrust and Basin-Range structure in southern Nevada, southwest Utah and northwest Arizona [M.S. thesis]: Corvallis, Oregon State University, 145 p. Carr, T.D., and Williamson, T.E., 2000, A review of Tyrannosauridae (Dinosauria: Coelurosauria) from New Mexico: New Mexico Museum of Natural History and Science Bulletin, v. 17, p. 113–145. Carr, T.D., and Williamson, T.E., 2005, Diversity of late Maastrichtian Tyrannosauridae (Dinosauria: Theropoda) from western North America: Journal of Vertebrate Paleontology (in press). Castle, J.W., Molz, F.J., Lu, S., and Dinwiddie, C.L., 2004, Sedimentology and fractal-based analysis of permeability data, John Henry Member, Straight Cliffs Formation (Upper Cretaceous), Utah, U.S.A.: Journal of Sedimentary Research, v. 74, no. 2, p. 270–284. Christensen, A.E., 2005, Sequence stratigraphy, sedimentology and provenance of the Drip Tank Member of the Straight Cliffs Formation, Kaiparowits Formation, Southern Utah [M.S. thesis]: Las Cruces, New Mexico State University, 96 p. Cifelli, R.L., 1990a, Cretaceous mammals of southern Utah. I. Marsupial mammals from the Kaiparowits Formation (Judithian): Journal of Vertebrate Paleontology, v. 10, p. 295–319. Cifelli, R.L., 1990b, Cretaceous mammals of southern Utah. II. Marsupials and marsupial-like mammals from the Wahweap Formation (early Campanian): Journal of Vertebrate Paleontology, v. 10, p. 320–331. Cifelli, R.L., 1990c, Cretaceous mammals of southern Utah. III. Therian mammals from the Turonian (early Late Cretaceous): Journal of Vertebrate Paleontology, v. 10, p. 332–345.

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Cifelli, R.L., 1990d, Cretaceous mammals of southern Utah. IV. Eutherian mammals from the Wahweap (Aquilan) and Kaiparowits (Judithian) formations: Journal of Vertebrate Paleontology, v. 10, p. 346–360. Cifelli, R.L., and Eaton, J.G., 1987, Marsupial from the earliest Late Cretaceous of Western U.S: Nature, v. 325, p. 520–522, doi: 10.1038/325520a0. Cobban, W.A., Dyman, T.S., Pollock, G.L., Takahashi, K.I., Davis, L.E., and Riggin, D.B., 2000, Inventory of dominantly marine and brackish-water fossils from Late Cretaceous rocks in and near Grand Staircase–Escalante National Monument, Utah, in Sprinkel, D.A., Chidsey, T.C., Jr., and Anderson, P.B., eds., Geology of Utah’s Parks and Monuments: Utah Geological Association Publication 28 (1st edition), p. 579–589. Currie, B.S., 1997, Sequence stratigraphy of nonmarine Jurassic-Cretaceous rocks, central Cordilleran foreland-basin system: Geological Society of America Bulletin, v. 109, p. 1206–1222, doi: 10.1130/00167606(1997)109<1206:SSONJC>2.3.CO;2. DeCelles, P.G., 2004, Late Jurassic to Eocene evolution of the Cordilleran thrust belt and foreland basin system, western U.S.A.: American Journal of Science, v. 304, p. 105–168. DeCourten, F.L., 1978, Nonmarine flora and fauna from the Kaiparowits Formation (Upper Cretaceous) of the Paria River Amphitheater, southwestern Utah: Geological Society of America Abstracts with Programs, v. 10, no. 3, p. 102. DeCourten, F.L., and Russell, D.A., 1985, A specimen of Ornithomimus velox (Therapoda, Ornithomimidae) from the terminal Cretaceous Kaiparowits Formation of southern Utah: Journal of Paleontology, v. 59, p. 1091–1099. Doelling, H.H., and Graham, R.L., 1972, Southwestern Utah coal fields: Alton, Kaiparowits, Plateau and Kolob-Harmony: Utah Geological and Mineral Survey, Monograph Series 1, 332 p. Doelling, H.H., and Davis, D.D., 1989, The geology of Kane County, Utah: geology, mineral resources, geologic hazards: Utah Geological and Mineral Survey Bulletin 124, 192 p. Doelling, H.H., Blackett, R.E., Hamblin, A.H., Powell, J.D., and Pollock, G.L., 2003, Geology of Grand Staircase–Escalante National Monument, Utah, in Sprinkel, D.A., Chidsey, T.C., Jr., and Anderson, P.B., eds., Geology of Utah’s Parks and Monuments: Utah Geological Association Publication 28 (2nd edition), p. 189–231. Dyman, T.S., Cobban, W.A., Titus, A.L., Obradovich, J.D., Davis, L.E., Eves, R.L., Pollock, G.L., Takahashi, K.I., and Hester, T.C., 2002, New biostratigraphic and radiometric ages for Albian-Turonian Dakota Formation and Tropic Shale at Grand Staircase–Escalante National Monument and Iron Springs Formation near Cedar City, Parowan, and Gunlock in SW Utah: Geological Society of America Abstracts with Programs, v. 34, no. 4, p. 13. Eaton, J.G., 1991, Biostratigraphic framework for Upper Cretaceous rocks of the Kaiparowits Plateau, southern Utah, in Nations, J.D., and Eaton, J.G., eds., Stratigraphy, depositional environments, and sedimentary tectonics of the western margin, Cretaceous Western Interior Seaway: Geological Society of America Special Paper 260, p. 47–63. Eaton, J.G., 1993, Therian mammals from the Cenomanian (Upper Cretaceous) Dakota Formation, southwestern Utah: Journal of Vertebrate Paleontology, v. 13, p. 105–124. Eaton, J.G., 1995, Cenomanian and Turonian (early Late Cretaceous) multituberculate mammals from southwestern Utah: Journal of Vertebrate Paleontology, v. 15, p. 761–784. Eaton, J.G., 1999, Vertebrate paleontology of the Iron Springs Formation, Upper Cretaceous, southwestern Utah, in Gillette, D.D., ed., Vertebrate Paleontology in Utah: Utah Geological Survey Miscellaneous Publication 99-1, p. 339–343. Eaton, J.G., 2002, Multituberculate mammals from the Wahweap (Campanian, Aquilan) and Kaiparowits (Campanian, Judithian) formations, within and near the Grand Staircase–Escalante National Monument, southern Utah: Utah Geological Survey Miscellaneous Publication 02-4, 66 p. Eaton, J.G., and Cifelli, R.L., 1997, Cretaceous vertebrates of the Grand Staircase–Escalante National Monument, in Hill, L.M., ed., Learning from the Land—Grand Staircase Escalante National Monument Science Symposium Proceedings: U.S. Department of the Interior, Bureau of Land Management, p. 365–372. Eaton, J.G., and Nations, J.D., 1991, Introduction; tectonic setting along the margin of the Cretaceous Western Interior Seaway, southwestern Utah and northern Arizona, in Nations, J.D., and Eaton, J.G., eds., Stratigraphy, depositional environments, and sedimentary tectonics of the western margin, Cretaceous Western Interior Seaway: Geological Society of America Special Paper 260, p. 1–8.

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Geological Society of America Field Guide 6 2005

Transect across the northern Walker Lane, northwest Nevada and northeast California: An incipient transform fault along the Pacific–North American plate boundary James E. Faulds Christopher D. Henry Nicholas H. Hinz Peter S. Drakos Benjamin Delwiche Nevada Bureau of Mines and Geology, MS 178, University of Nevada, Reno, NV 89557, USA

ABSTRACT Within the western Great Basin, a system of dextral strike-slip faults accommodates a significant fraction of the North American–Pacific plate motion. The northern Walker Lane in northwest Nevada and northeast California occupies the northern terminus of this fault system and is one of the youngest and least developed parts of the North American–Pacific transform plate boundary. Accordingly, the northern Walker Lane affords an opportunity to analyze the incipient development of a major strike-slip fault system. In northwest Nevada, the northern Walker Lane consists of a discrete ~50-km-wide belt of overlapping, curiously left-stepping dextral faults, whereas a much broader zone of disconnected, widely-spaced northwest-striking faults characterizes northeast California. The left steps accommodate little shortening and are not typical restraining bends. The left-stepping dextral faults may represent macroscopic Riedel shears developing above a nascent lithospheric-scale transform fault. Strands of the northern Walker Lane terminate in arrays of northerly striking normal faults in the northwestern Great Basin and along the eastern front of the Sierra Nevada. These relations suggest that dextral shear in the northern Walker Lane is transferred to ~NW-SE extension in the Great Basin. Offset segments of a west-trending Oligocene paleovalley suggest ~20–30 km of cumulative dextral slip across the northern Walker Lane. Strike-slip faulting began between 3 and 9 Ma, indicating a long-term slip rate of ~2–10 mm/yr, which is compatible with GPS geodetic observations of the current strain field. Keywords: Walker Lane, Nevada, paleovalley, Riedel shear, transform fault. INTRODUCTION The western margin of North America contains a broad zone of distributed shear stretching from the San Andreas fault system to the Basin and Range province (Fig. 1; Wernicke, 1992;

Atwater and Stock, 1998). GPS geodetic results indicate that a system of right-lateral strike-slip faults in the western Great Basin, known as the Walker Lane in its northern reaches (Locke et al., 1940; Stewart, 1988) and the eastern California shear zone to the south (Dokka and Travis, 1990), accommodates as much

Faulds, J.E., Henry, C.D., Hinz, N.H., Drakos, P.S., and Delwiche, B., 2005, Transect across the northern Walker Lane, northwest Nevada and northeast California: An incipient transform fault along the Pacific–North American plate boundary, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 129–150, doi: 10.1130/2005.fld006(06). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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A

B

C

Figure 1. Cenozoic tectonic evolution, western North America (modified from Atwater and Stock, 1998). (A) 30 Ma. (B) 10 Ma. (C) 0 Ma. The San Andreas fault system has progressively lengthened over the past 30 Ma, as more of the Pacific plate has come into contact with North America. The Walker Lane currently accommodates 10%–25% of the Pacific–North American plate motion. The box in (C) surrounds the northern Walker Lane and adjacent parts of the Great Basin. ACA—ancestral Cascade arc; CA—Cascade arc; MTJ—Mendocino triple junction; SAF—San Andreas fault.

Figure 2. Faults of the Walker Lane and San Andreas system (modified from Stewart, 1988) on shaded elevation model. Circles in central Nevada are calderas, including sources of ash-flow tuffs of this trip. Box shows area of Figure 5. P—Poco Canyon caldera, source of tuff of Chimney Spring (John, 1995); E—Elevenmile Canyon caldera, source of tuff of Painted Hills; C—Campbell Creek caldera, source of tuff of Campbell Creek; OV—Owens Valley.

as 20%–25% of dextral motion between the North American and Pacific plates (Thatcher et al., 1999; Dixon et al., 2000; Oldow et al., 2001; Bennett et al., 2003; Hammond and Thatcher, 2004). The Walker Lane essentially accommodates dextral motion of the Sierra Nevada block relative to the Great Basin and marks an abrupt physiographic change, whereby the predominant northnortheast–trending topographic grain in the Great Basin gives way westward to more heterogeneous terrain (Fig. 2). To the south, this system of faults merges with the San Andreas fault in southern California, whereas to the north it terminates in northeast California near the southern end of the Cascade arc. Considering geodetic data showing that the Walker Lane accommodates a significant fraction of the North American– Pacific plate motion, as well as the progressive northward migration of the Mendocino triple junction and related lengthening of the San Andreas fault (e.g., Atwater and Stock, 1998), it follows that the Walker Lane may also have propagated northwestward through the western Great Basin. Estimates of cumulative dextral offset across the eastern California shear zone and Walker Lane since Miocene time are ~50–100 km in eastern California (e.g., Dokka and Travis, 1990), 48–75 km in west-central Nevada (Ekren and Byers, 1984; Oldow, 1992), and essentially zero at its northwestern terminus in northeastern California. Thus, the northern Walker Lane is one of the least developed and possibly youngest parts of the transform boundary. Analysis of an incipient strike-slip fault system is therefore possible in this region. The purpose of this field trip is to (1) view the general geometry of the northern Walker Lane, (2) use an approximately west-trending system of Oligocene paleovalleys to constrain

Transect across the northern Walker Lane

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Figure 3. Map showing field trip route for each day and major physiographic features in the northern Walker Lane.

cumulative dextral slip across the northern Walker Lane, and (3) evaluate kinematic models for this youthful strike-slip fault system. The trip will traverse from east to west across the northern Walker Lane just north of Reno, Nevada (Fig. 3). This work incorporates recently completed detailed geologic mapping, structural analysis, and geochronologic and paleomagnetic studies in the region (Cashman and Fontaine, 2000; Faulds and Henry, 2002; Henry et al., 2003; 2004a, 2004b, 2006; Faulds et al., 2003a, 2003b, 2004b, 2005b; Hinz, 2004).

GEOLOGIC SETTING As western North America has evolved from a convergent to a transform margin in the past 30 m.y., the northern Walker Lane and surrounding regions have experienced widespread volcanism and tectonism. Tertiary volcanic strata include ca. 31–23 Ma ashflow tuffs associated with the southward migrating “ignimbrite flare-up,” ca. 22–5 Ma calc-alkaline intermediate rocks related to the ancestral Cascade arc, and ca. 13 Ma to present bimodal

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rocks linked to Basin and Range extension (e.g., Deino, 1985; Stewart, 1988; Best et al., 1989; Christiansen and Yeats, 1992; Garside et al., 2000; Henry et al., 2004a; Henry and Faulds, 2004). In the past 13–3 m.y., coincident with the northward migration of the Mendocino triple junction and associated termination of subduction, arc volcanism retreated northwestward, Basin and Range normal faulting advanced westward into the Sierra Nevada (Dilles and Gans, 1995; Henry and Perkins, 2001; Surpless et al., 2002), and strike-slip faulting began in the northern Walker Lane. East-west extension produced north-trending basins, within which accumulated thick sedimentary sections (Trexler et al., 2000). Tertiary strata rest directly on Mesozoic granitic and metamorphic basement. The ignimbrite flare-up is particularly important to unraveling the evolution of the northern Walker Lane, because thick sequences of ash-flow tuffs were deposited in paleovalleys that extended across western Nevada and California (e.g., Lindgren, 1911; Henry et al., 2003; Faulds et al., 2005b). The paleovalleys were part of an extensive Oligocene (Eocene to Miocene?) paleodrainage system that extended westward from an orogenicvolcanic highland in what is now the central Great Basin to the Pacific Ocean, which was in the Great Valley at the time (Christiansen and Yeats, 1992; Dilek and Moores, 1999; Henry et al., 2003; Faulds et al., 2005b). The Oligocene paleogeography was

quite different than that of today. Continuity of the paleovalleys across what is now the Sierra Nevada–Basin and Range boundary demonstrates that the Sierra Nevada was topographically lower than the Great Basin at the time (our work), even though the Sierra Nevada may have stood at higher elevations than it does today (Wernicke et al., 1996). Paleovalleys also cross the Walker Lane at nearly right angles, thereby providing piercing lines with which to estimate offset on dextral faults. The paleovalleys contain distinctive 31–23 Ma ignimbrite sections (Fig. 4), which reflect the southward progression of magmatism. The southward migration of volcanism is critical, because ash-flow tuffs filling paleovalleys in the northern Walker Lane are not easily confused with generally younger sections farther south. The ash flows erupted from calderas in central Nevada, well east of the Walker Lane (Best et al., 1989; John, 1995; Henry et al., 1997). In western Nevada and northeast California, this volcanism terminated at ca. 23 Ma, before development of the Walker Lane. Thus, offset Oligocene paleovalleys record cumulative displacement across the northern Walker Lane (Fig. 5). Arc-related volcanism immediately followed the ignimbrite flare-up in much of western Nevada, thus preserving many tuff-filled paleovalleys beneath a cover of younger resistant volcanic rock. For the purpose of this field trip, the northern Walker Lane includes the system of right-lateral faults within the western

Figure 4. (A) Table showing ash-flow tuffs in paleovalley segments. Shaded box indicates presence of tuff in paleovalley at that location. (B) Schematic ash-flow tuff stratigraphy in paleovalley at Dogskin Mountain, with ages corresponding to individual tuffs in (A). Paleovalley is ~8 km wide, and ash-flow sequence is as much as 700 m thick.

Transect across the northern Walker Lane Great Basin from about the latitude of Carson City northwestward to the southern part of the Cascade arc in the Modoc Plateau region of northern California. The northern Walker Lane consists of a complex system of kinematically linked northwest-striking, left-stepping right-lateral faults, northerly striking normal faults, and subordinate E-NE–striking sinistral faults (e.g., Stewart, 1988; Cashman and Fontaine, 2000; Faulds et al., 2005b; Henry et al., 2006). From east to west, major dextral faults include the Pyramid Lake, Warm Springs Valley, Honey Lake, and Mohawk Valley fault zones (Fig. 5). In northeast California, the discrete belt of strike-slip faults gives way to a diffuse zone of widelyspaced northwest-striking dextral faults and lineaments that extends northwestward to the southern part of the Cascade arc (Fig. 2; Grose, 2000; Colie et al., 2002).

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OLIGOCENE PALEOVALLEYS: PIERCING LINES ACROSS THE NORTHERN WALKER LANE Well-defined paleovalley axes and/or margins within the Oligocene drainage network, stretching from the Nightingale Mountains in northwest Nevada to the Diamond Mountains in northeast California, provide piercing lines with which to estimate offset on dextral faults in the northern Walker Lane (Fig. 5). Petrographic analysis and 40Ar/39Ar geochronology demonstrate that all segments of this paleovalley system, including at least two tributaries, contain an essentially identical, albeit westwardthinning sequence of 31.3–23.5 Ma ash-flow tuffs (Fig. 4). This includes several distinctive and widespread units, including the 25.3 Ma Nine Hill Tuff (c.f., Deino, 1985), the 28.8 Ma tuff of

Figure 5. Northern Walker Lane showing left-stepping, NW-striking dextral faults and offset paleovalley axes (modified slightly from Faulds et al., 2005b). Paleovalleys are grouped into three sets. Individual segments in each set are inferred to represent an originally continuous paleovalley. DM—Diamond Mountains; DS—Dogskin Mountain; FR—Fox Range; FS—Fort Sage Mountains; HLF—Honey Lake fault; HP—Haskell Peak; LR—Lake Range; MVF—Mohawk Valley fault; NR—Nightingale Range; PLF—Pyramid Lake fault; PM—Peterson Mountain; PR—Pah Rah Range; SL—Seven Lakes Mountain; SV—Sierra Valley; TH—Terraced Hills; VM—Virginia Mountains; WSF—Warm Springs Valley fault.

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Campbell Creek, and the 31.3 Ma tuff of Axehandle Canyon. Thin fluvial gravel deposits, commonly with large rounded boulders of tuff or basement rock, are sandwiched between some tuffs. Such gravels indicate deposition in a robust fluvial system in the westward flowing paleodrainage network. The famed “Auriferous Gravels” of the Sierra Nevada were deposited in downstream parts of the system (Lindgren, 1911; Garside et al., 2005). In the western part of the region, the Oligocene ash-flow tuff sequence thins significantly, concomitant with appreciable increases in thickness of interbedded fluvial units. In the northern Sierra Nevada and Fort Sage and Diamond Mountains, cumulative thicknesses of tuff range from ~200–250 m, which contrasts with 500–700 m in the Virginia and Nightingale Mountains. Western areas do, however, contain essentially the same sequences of tuffs, but individual tuffs are generally thinner and more tuffs are missing compared to areas farther east. The thinner sequences of tuffs did not result from relative uplift (relative to the Great Basin) and erosion of the Sierra Nevada and related blocks, because middle to late Miocene volcanic rocks, which predate such uplift at these latitudes (e.g., Trexler et al., 2000; Henry and Perkins, 2001), cap the late Oligocene ash-flows. The thinner sequences of tuffs in these areas probably mostly reflect greater distances from source calderas, as well as perhaps a westward transition to broader, less confined paleovalleys. Our mapping shows that the paleovalleys were relatively broad and shallow, in contrast to the narrow v-shaped valleys of the present Sierra Nevada. East of the Pyramid Lake fault, the Nightingale Mountains contain a well-defined, ~W-SW–trending, ~10 km wide paleo-

valley (Fig. 5, segment 1A) with a well-exposed abrupt northern margin (Fig. 6), along which the entire >700 m thick sequence of tuffs pinches out. Eleven ash-flow tuffs have been observed in this paleovalley segment. In the southern Lake Range, however, only a thin veneer (<30 m thick) of Oligocene tuff crops out, indicating that the westward continuation of the paleovalley projects south of the Lake Range. West of the Pyramid Lake fault in the Virginia Mountains, the northern part of a W-SW–trending paleovalley contains a 600 m thick sequence of 15 ash-flow tuffs (Fig. 5, segment 1B; Faulds et al., 2003a, 2003b). Because of nearly identical tuff sequences, the paleovalley in the Virginia Mountains is correlated with that in the Nightingale Mountains. Based on its westerly trend, this paleovalley has been offset ~5–10 km across the Pyramid Lake fault (Fig. 7). Farther south near Fernley, Nevada, an E-W–trending late Miocene anticline appears to be offset ~10 km across the Pyramid Lake fault. Similar relations indicate ~10 km of right-lateral displacement across the Warm Springs Valley fault. On the northwest flank of Dogskin Mountain, more than 500 m of 31.3–24.9 Ma ash-flow tuff (16 units) pinch out against a well-exposed southeast margin of a W-SW–trending trending paleovalley (Fig. 5, segment 1C, and Fig. 8) (Henry et al., 2004b). In addition to the same suite of ash-flow tuffs, a distinctive rock-avalanche megabreccia derived from collapse of nearby rhyolite domes is found directly above a 24.9 Ma ash-flow tuff in both the Virginia Mountains and at Dogskin Mountain. The distribution of ash-flow tuffs in the structural block between the Pyramid Lake and Warm Springs Valley faults

Figure 6. Abrupt northern margin (dashed line) of paleovalley in the Nightingale Mountains (looking west). The 31.3 Ma tuff of Axehandle Canyon crops out along the paleovalley margin, where primary compaction foliation in the tuff approaches 60°. Stop 1-2 traverses to the contact denoted by black arrow.

Transect across the northern Walker Lane (Virginia Mountains and Pah Rah Range) is much more extensive than that in the Nightingale Mountains or at Dogskin Mountain (Fig. 5, segment 1B and 2B). Although a small paleo-ridge of Mesozoic basement exists in the northern Pah Rah Range, the overall distribution of tuffs within this block suggests a 20–30km-wide paleovalley. We attribute these relationships to result from the confluence of two paleovalleys, with a northern ~WSW–trending trending branch (e.g., Nightingale Mountains) and a southern W-NW–trending tributary that skirts the northern margin of the Carson Sink. Due to thick Neogene basin fill within the Honey Lake basin, displacement across the Honey Lake fault is more difficult to

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constrain. However, nearly identical sequences of tuff crop out at Dogskin, Seven Lakes, and northern Peterson Mountains (Fig. 5, segments 1C and 1D). The southern margin of a paleovalley in the northern part of Peterson Mountain probably correlates with the paleovalley margin at Dogskin Mountain. This suggests 3–6 km of dextral offset along the southeastern part of the Honey Lake fault. An apparent northern tributary to this paleovalley is marked by very similar 200–250-m-thick sections of ash-flow tuff in the Fort Sage Mountains and north-central Diamond Mountains (Fig. 5, segments 3A and 3B). Although additional deposits of tuff may be obscured by the Honey Lake basin, the lack of similar sequences of tuff in other parts of the Diamond Mountains supports a corre-

Figure 7. Aerial view looking east at offset Oligocene paleovalleys across the Pyramid Lake fault (PLF), showing thick sequence of tuffs in axial parts of paleovalleys in the Virginia Mountains (foreground) and Nightingale Mountains (background).

Figure 8. Looking southwest at the southeast margin of 600-m-deep paleovalley in the northwestern part of Dogskin Mountain. The entire ash-flow tuff sequence (15 units) pinches out against this paleovalley margin.

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lation between these paleovalley segments. If correct, the central part of the Honey Lake fault accommodated ~10–15 km of dextral offset, assuming a westerly trending paleovalley (Hinz, 2004). As the Honey Lake fault zone dies out to the southeast, much of the right-lateral displacement may be taken up by major normal faults bounding Long Valley and folding and minor thrusting in a small restraining bend at Seven Lakes Mountain. The westward continuation of the main stem of the paleovalley system projects into the Sierra Valley area and is probably buried by Neogene sediments. A relatively thick sequence of correlative Oligocene ash-flow tuff crops out west of Sierra Valley at Haskell Peak (Brooks et al., 2003) in the northern Sierra Nevada, roughly on trend with this paleovalley segment. The westernmost strand of the northern Walker Lane, the Mohawk Valley fault, has accommodated only minor dextral offset. Due to scant exposures of Oligocene tuff, definitive paleovalley segments have not been observed in this area. However, an intrusive contact between Cretaceous batholithic and Paleozoic metamorphic rocks shows little offset across the fault zone (Fig. 5; Saucedo and Wagner, 1992), which suggests that the Mohawk Valley fault has accommodated negligible (<1 km) right-lateral displacement. In summary, inferred offsets of the Oligocene paleovalley system indicate ~20–30 km of cumulative dextral displacement, as measured orthogonal to the northern Walker Lane from the southern part of Pyramid Lake on the northeast to the southeast end of the Honey Lake fault on the southwest (Fig. 5). Because of the en echelon fault pattern, the estimate of total displacement includes offset from only overlapping parts of the strike-slip fault system, as measured orthogonal to the overall trend of the northern Walker Lane. Depending on where the cross section line is drawn, this may include minor slip near the ends of some strike-slip faults and nearly maximum displacement on other faults. TIMING AND RATES OF STRIKE-SLIP FAULTING Several key relationships suggest a relatively recent onset of strike-slip deformation in the northern Walker Lane. For example, tilts are generally concordant both within the sections of Oligocene ash-flow tuff and between middle Miocene volcanic rocks and Oligocene tuffs (Fig. 9), suggesting that no appreciable deformation accompanied the ignimbrite flare-up and ancestral Cascade arc volcanism. Northwest-striking ca. 22 Ma veins in Pah Rah Range, previously attributed to early Walker Lane deformation (Wallace, 1975), probably record minor NE-SW extension, an unlikely consequence of NW-trending dextral shear (Faulds et al., 2005a). Although ~E-W extension initiated basin development ca. 13 Ma in the region (Trexler et al., 2000), there is no evidence that strike-slip faulting accompanied early stages of Miocene extension. Moreover, it is important that 3.5 Ma sediments along the Warm Springs Valley fault are as highly deformed as the Oligocene tuffs (Henry et al., 2006). Furthermore, ~25° of clockwise vertical-axis rotation within the Walker Lane east of

Figure 9. East-tilted Oligocene ash-flow tuffs and middle Miocene Pyramid sequence in the Virginia Mountains just north of Sutcliffe, Nevada (near Stop 1-5). Note that tilts in the Oligocene tuffs (Twc) and middle Miocene mafic lavas (Tpb) are approximately concordant (dashed lines indicate layering orientation). The sequence of tuffs pinches out rapidly northward in this area along the northern margin of the Virginia Mountains paleovalley. Twc—30.1 Ma tuff of Cove Spring; Tpb—15–13 Ma mafic lavas of the Pyramid sequence; Kgr—Cretaceous granite. Stop 1-5 is located near the Kgr label.

Carson City is bracketed between ca. 9 and 5 Ma (Cashman and Fontaine, 2000). Considering these relations, we infer that significant strike-slip faulting began ca. 9–3 Ma. Such timing coincides with the northward migration of the Mendocino triple junction into these latitudes (Atwater and Stock, 1998). Recent movement on strike-slip and normal faults has been broadly coeval in the northern Walker Lane, as evidenced by Quaternary fault scarps and historical seismicity (e.g., Bell, 1984; dePolo et al., 1997; Ichinose et al., 1998). The inferred magnitude (20–30 km of dextral slip across the northern Walker Lane) and timing of deformation suggest long-term slip rates of ~2–10 mm/yr, which is compatible with published geodetic data from the region (e.g., Bennett et al., 2003; Hammond and Thatcher, 2004). KINEMATIC MODEL The en echelon pattern of dextral faults, minimal offset (5– 15 km) on individual faults, and relatively small cumulative displacement indicate that the northern Walker Lane is an incipient strike-slip fault system. Although small left steps on individual faults have induced local shortening, the broad left steps between major faults accommodate little, if any, shortening and are unlike typical restraining bends. Hard linkages between some major strands of the northern Walker Lane (e.g., Pyramid Lake and Warm Springs Valley faults) appear to be absent. Left-stepping dextral faults also characterize the central part of the Walker Lane in west-central Nevada (see Oldow, 1992). The left-stepping, en echelon geometry resembles patterns of primary Riedel shears (cf., Petit, 1987) developed above strike-slip faults in clay models

Transect across the northern Walker Lane

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B

Figure 10. (A) Riedel shear model for northern Walker Lane and analogous clay model from Wilcox et al. (1973). Coeval NW-directed dextral shear and W-NW extension may account for slight counterclockwise rotation of fault blocks that collapse in domino-like fashion to accommodate extension. (B) Schematic block diagram of incipient strike-slip fault system with primary Riedel shears developing above a through-going fault at depth.

(Fig. 10; e.g., Wilcox et al., 1973). We therefore conclude that the dextral faults within the northern Walker Lane are primary Riedel shears developing above a nascent lithospheric-scale transform fault within the upper mantle and/or lower crust (Faulds et al., 2003c, 2005b). The E-NE–striking sinistral-normal faults may be analogous to secondary Riedel shears. On the basis of the ~N40°W alignment of major strike-slip faults in the northern Walker Lane, the presumed through-going master fault zone beneath the northern Walker Lane may strike ~N50°W. This incipient transform fault may represent a reactivated Mesozoic strike-slip fault (e.g., Oldow, 1984; Henry et al., 2006), which originally accommodated oblique convergence between the Pacific and Farallon plates. Alternatively, dextral shear associated with the current transform margin may have induced development of a new lithospheric-scale structure. Major dextral faults in the northern Walker Lane terminate in arrays of northerly striking normal faults, suggesting that dextral shear in the northern Walker Lane is transferred to W-NW– directed extension in the northwest Great Basin and along the east front of the Sierra Nevada. In the western part of the northern Walker Lane, major strike-slip faults merge southward with a system of east-dipping normal faults, many of which coalesce southward to form the Sierra Nevada frontal fault system. To the east, dextral shear along the Pyramid Lake fault appears to bleed off into major west-dipping range-front faults along the Lake Range and Nightingale Mountains. Thus, each en echelon strikeslip fault within the northern Walker Lane appears to terminate in broad extensional domains. Such relationships imply that much of the extension in the northwestern Great Basin and along the east side of the Sierra Nevada is accommodating the northwestern terminus of dextral shear within the Walker Lane. Because all of the dextral shear within the Walker Lane is ultimately diffused into extension within the northwestern Great Basin, the Sierra Nevada block and northwesternmost part of the Great Basin are essentially moving northwestward as one relatively coherent block (Fig. 11; Faulds et al., 2004a).

In northeast California, the diffuse zone of faulting at the northwest terminus of the Walker Lane mimics clay model experiments involving broadly distributed shear with no through-going fault at depth (e.g., An and Sammis, 1996). In this early stage of strike-slip faulting, the minor faults are also primary Riedel shears, but the strain is accommodated across a much broader zone due to the lack of a through-going fault at depth. Evidence for this earlier

Figure 11. General kinematics of the northwestern Great Basin (from Faulds et al., 2004a). The Walker Lane accommodates northwestward translation of the Sierra Nevada block relative to the central and eastern Great Basin. As the Walker Lane terminates northwestward, dextral motion is diffused into NW-directed extension along NNE-striking normal faults within the northwestern Great Basin. With little relative motion between them, the northern part of the Sierra Nevada block and northwestern-most part of the Great Basin are essentially a coherent structural block.

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stage abounds in northwest Nevada, as the blocks between the major active strike-slip faults commonly contain minor inactive NW-striking dextral faults (Faulds et al., 2003a, 2003b). Thus, the NW-propagating northern Walker Lane appears to record, in the along-strike dimension, two initial stages in the development of a major intracontinental strike-slip fault system. In northeast California, where dextral motion is minimal, a broad zone of disconnected, en echelon faults has developed. In northwest Nevada, where strike-slip faulting began between ca. 9 and 3 Ma and ~20–30 km of slip has accumulated, deformation has been concentrated along a much narrower zone of en echelon faults, or possible Riedel shears, presumably above a throughgoing strike-slip fault in the upper mantle and/or lower crust. The modest (~15–25°) counterclockwise rotation of three major fault blocks in the northern Walker Lane (Cashman and Fontaine, 2000; Faulds et al., 2004b) is opposite the typical clockwise rotation in dextral shear zones. In the transtensional setting of the northern Walker Lane, coeval dextral shear in the left-stepping fault system and W-NW regional extension may account for the counterintuitive anticlockwise rotation (Fig. 10). This sense of rotation would ultimately rotate Riedel shears toward the orientation of the main shear zone at depth, thus facilitating development of a through-going, upper-crustal strike-slip fault (Faulds et al., 2005b). Ironically, as such a fault develops and the strike-slip fault system matures, the sense of vertical-axis rotation may reverse and become compatible with the overall sense of shear strain. The apparent change from anticlockwise to clockwise rotations southeastward within the northern Walker Lane (Cashman and Fontaine, 2000) may reflect such a progression. Such complex kinematics may characterize the early evolution of strike-slip fault systems in both transtensional and transpressional settings. The kinematics of the northern Walker Lane may therefore be applicable to other propagating intracontinental strike-slip fault systems, such as the Anatolian fault (e.g., Taymaz et al., 1991; Armijo et al., 2004) and some strike-slip fault systems in the Himalayan orogenic belt (e.g., Baljinnyam et al., 1993; Jackson et al., 1995; Bayasgalan et al., 1999). If the Mendocino triple junction eventually migrates northward to the Oregon coast, the Walker Lane may ultimately afford a more stable configuration for the Pacific–North America plate boundary. It is also important to note that the San Andreas fault system has a history of stepping inland with time (Atwater and Stock, 1998). The Walker Lane may therefore reflect the birth of a lithospheric-scale transform fault and presage an eventual eastward jump in the plate boundary. DAY 1—NIGHTINGALE TO VIRGINIA MOUNTAINS ACROSS PYRAMID LAKE FAULT Introduction Figure 3 shows an overview of the trip route for each day. The trip will traverse from east to west across the northern Walker Lane at ~40°N latitude just north of Reno, Nevada.

On Day 1, we will begin tracking a W-SW–trending Oligocene paleovalley system from just east of the northern Walker Lane in the Nightingale Mountains westward across the rightlateral Pyramid Lake fault to the Virginia Mountains within the northern Walker Lane (Fig. 5, segments 1A and 1B). This paleovalley serves as a piercing line with which to measure offset across right-lateral faults. We will also discuss the geometry and kinematics of major strike-slip faults in the northern Walker Lane and possible kinematic models for this incipient strikeslip fault system. En Route Discussion As we leave Reno and travel east on I-80 (Fig. 3), we travel through the Truckee River Canyon, which is primarily carved into thick middle to late Miocene volcanic rocks associated with both the ancestral Cascade arc and extension-related bimodal magmatism. The canyon parallels the E-NE–striking, left-lateral Olinghouse fault, which may have ruptured in a 6.7 magnitude earthquake in 1869, with as much as 3.6 m of left slip (Sanders and Slemmons, 1979; dePolo et al., 1997). Rhomb-shaped depressions in left stepovers and offset of stream channels and debris flow levees indicate a left-lateral component. Strands of the Olinghouse fault zone straddle I-80 from about Exit 23 eastward to the mouth of the Truckee River Canyon near Exit 43. Directions to Stop 1-1 From the University of Nevada–Reno (UNR) campus, head east on Interstate 80 (I-80) for ~30 mi, then take Exit 43 to the town of Wadsworth. After exiting I-80, turn left onto Hwy 427 and follow 1.3 mi to Wadsworth, then turn left onto Hwy 447. About 9 mi to the north, Hwy 447 crosses the trace of the Pyramid Lake fault. At ~10 mi north on Hwy 447, note Lake Lahontan sedimentary rocks in Truckee River Canyon to the east. At 11.5 mi, Pyramid Lake and Anaho Island can be observed to the north. After traveling north for 12.5 mi on Hwy 447, turn left onto the dirt road and proceed through gate. After 0.25 mi veer left and continue 1.4 mi to Stop 1-1. We take a short hike on the ridge to the east of the road. Stop 1-1: Pyramid Lake Fault The ridge provides an excellent view of a small pull-apart along the right-lateral Pyramid Lake fault (Fig. 12), as well as surrounding mountain ranges. To the west, well exposed but altered Oligocene ash-flow tuffs fill a paleovalley in the Pah Rah Range (Fig. 5, segment 2B). These units are best exposed in the southern Pah Rah Range, where they comprise the colorful bands of rock as best seen just north of Wadsworth. To the east, a thick sequence of middle to late Miocene basalt flows dominates the Truckee Range. The Pyramid Lake fault lies directly to our west at the foot of the ridge. It is the easternmost of four major left-stepping, NW-striking dextral faults in the northern Walker Lane (Fig. 5). The broad left steps between the faults are not typical restraining bends, as evidenced by the

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lack of shortening within the intervening fault blocks. The en echelon left-stepping pattern may instead reflect primary Riedel shears developing above a through-going fault within the lower crust and/or upper mantle (Fig. 10; Faulds et al., 2005b). Considering the tendency of the San Andreas fault system to step inland through time (e.g., Atwater and Stock, 1998), this underlying through-going fault may be an incipient transform fault. All major strike-slip faults within the northern Walker Lane have Holocene scarps along much of their lengths. Based on recent trenching and the apparent offset of drainages near this stop, Briggs and Wesnousky (2004) concluded that the Pyramid Lake fault has ruptured in four major earthquakes in the Holocene, with inferred slip rates of 2.6 ± 0.3 mm/yr. However, Bell and House (2005) suggested that some strands of the Pyramid Lake fault have accommodated primarily normal offset and that slip rates may therefore be significantly less. Directions to Stop 1-2 From Stop 1, return to Hwy 447, turn left and proceed 3.8 mi through the town of Nixon. Just north of Nixon, turn right onto a dirt road and then immediately (~100 ft) turn left onto a graded dirt road. At ~6.9 mi on the graded road, a good view can be seen of a thick sequence of Oligocene ash-flow tuffs filling a paleovalley in the Nightingale Mountains to the northeast. At 9.1 mi, continue straight at road intersection and ignore road closed sign. At ~12.3 mi, turn right into a canyon and proceed 0.8 mi to Stop 1-2, where we will take a short but steep hike to the crest of the ridge to the west of the road. Stop 1-2: Nightingale Mountains Paleovalley Margin The crest of this ridge provides a spectacular view of the northern margin of a large W-SW–trending Oligocene paleovalley (Fig. 6). The Oligocene paleovalleys in this region provide piercing lines with which to gauge offset along the major strikeslip faults. More than 700 m of ash-flow tuff, including at least 11 distinct units, were deposited in this paleovalley. The entire 700 m thick section of tuffs pinches out against the northern margin. In this area, we see the three lower tuffs in the section, including the 31.3 Ma tuff of Axehandle Canyon, the 31.2 Ma tuff of Hardscrabble Canyon, and the 31.0 Ma tuff of Rattlesnake Canyon (Faulds et al., 2005b). We have thus far identified 18 distinct ashflow tuffs within the Oligocene paleovalleys in this region (Fig. 4). These tuffs were derived from calderas in central Nevada (Fig. 2). On our traverse, we will examine the tuff of Axehandle Canyon, with compaction foliations approaching 60° along the steep erosional escarpment of the paleovalley. The steep dips of the compaction foliation are a primary feature and flatten out within a few hundred meters of the paleovalley margin. Here, the paleovalley is cut into Jurassic (?) metasedimentary rocks. Abundant lithic fragments of the metasedimentary rock are found in the tuffs along the paleovalley margin. On several subsequent stops (e.g., Stop 6 on Day 1 and Stop 6 on Day 3), we will view the same ash-flow tuffs in more western parts of the paleovalley system.

Figure 12. View looking northwest along Pyramid Lake fault with Pyramid Lake in background. A small pull-apart basin (marked by the small playa) has developed in a small right step along the fault and is near Stop 1-1.

Directions to Stop 1-3 From Stop 2, return to Hwy 447, turn right (north), continue 7.1 mi, and turn right onto a short dirt road to the top of small ridge. Stop 1-3: Overview of Lake Range and Nightingale Mountains This ridge provides an excellent view of the entire paleovalley in the Nightingale Mountains to the east, as well as gently east-tilted middle Miocene mafic to intermediate volcanic rocks in the southern Lake Range to the northwest. The 10 km wide W-SW–trending paleovalley in the Nightingale Mountains projects westward toward the south end of Pyramid Lake. Only a thin veneer of Oligocene ash-flow tuff is found in the Lake Range to the northwest. The southern Lake Range is dominated by mafic lavas of the middle Miocene Pyramid sequence (Bonham and Papke, 1969). Voluminous ancestral Cascade arc volcanism in the region helped to preserve many of

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the tuff-filled paleovalley sequences beneath a cap of resistant rock. The porphyritic basalt that crops out on this hill is part of the Pyramid sequence. This stop also provides an overview of kinematic relations between major strike-slip and normal faults within the region. Similar to other major strike-slip faults in the northern Walker Lane, the Pyramid Lake fault terminates in complex arrays of normal faults. In this area, slip on the Pyramid Lake fault appears to be transferred to west-dipping range-front normal faults that bound the Lake Range and Nightingale Mountains on the west. These range-front faults merge southward with the Pyramid Lake fault (Fig. 5). Directions to Stop 1-4 From Stop 3, return to Hwy 447, turn right, and continue 0.6 mi to north, then turn left onto a graded dirt road. Continue 8.3 mi on the graded road, then turn left just east of the Pyramid. Keep to right for 0.5 mi, then veer left to Stop 1-4 near the shoreline of Pyramid Lake. Stop 1-4: Pyramid Lake and Southern Lake Range At this stop, we have excellent views of both Pyramid Lake and the lower part of the Tertiary section in the southern Lake Range. Pyramid Lake is fed by the Truckee River, which emanates from Lake Tahoe and is supplied largely by snow melt in the Sierra Nevada. Pyramid Lake is one of the last vestiges of Lake Lahontan, which inundated many of the basins in western Nevada and northeast California during its highstand in the late Pleistocene (e.g., Adams et al., 1999). The Pyramid is the small but conspicuous pyramidal-shaped island located just offshore to the west. It is one of many large tufa mounds that grace the landscape surrounding Pyramid Lake. Tufa is a calcium carbonate

deposit that formed as carbonate-rich springs discharged into Lake Lahontan (Benson, 1994). A short hike from the road takes us to an exposure of the Oligocene nonconformity, where a thin (10–15-m-thick) deposit of the 31.3 Ma tuff of Axehandle Canyon rests on Cretaceous granite (Fig. 13). The presence of only a thin veneer of Oligocene tuffs throughout the Lake Range indicates that the paleovalley observed in the Nightingale Mountains must project south of the Lake Range toward the southern end of Pyramid Lake. Thick sequences (>1 km) of gently east-tilted middle Miocene (ca. 15–13 Ma) basaltic andesite lavas dominate the high ridges to the east. Tilt magnitudes are similar within the entire Tertiary section of the Lake Range, suggesting that major extension occurred after ca. 13 Ma. A major west-dipping range-front normal fault bounds the Lake Range on the west. The Pyramid Lake basin is a large easttilted half graben in the hanging wall of this fault. The right-lateral Pyramid Lake fault projects into the southern end of Pyramid Lake but does not appear to exit the north end of the ~40-km-long lake, at least not as a discrete fault. This suggests that slip on the Pyramid Lake fault is transferred to the west-dipping normal fault system that includes range-front faults along the west flanks of the Lake Range, Fox Range, and Nightingale Mountains. Minor strands of the Pyramid Lake fault may, however, project through the Terraced Hills directly northwest of the lake (Fig. 5) and connect with NNEstriking normal faults in the Smoke Creek Desert area. Directions to Stop 1-5 Return to Hwy 447 and turn right. Continue 9.0 mi to the south traveling back through Nixon, then turn right onto Hwy 446 just south of Nixon, which heads northwest along the west shore of Pyramid Lake. After 13.4 mi, turn right onto Hwy 445 at stop sign and proceed north passing through the town of Sutcliffe. At 4.6 mi past the Hwy 446-445 intersection, turn left onto a dirt

Figure 13. Oligocene tuffs in the Lake Range. Only a thin veneer of Oligocene ash-flow tuffs crop out in the Lake Range. Here, a 10–15-m-thick section of the 31.3 Ma tuff of Axehandle Canyon rests on Cretaceous granite (Stop 14). Geologist for scale is enclosed by black circle.

Transect across the northern Walker Lane road and pass through a gate. Continue west on dirt road for 0.6 mi to Stop 1-5. Stop 1-5: Northern Margin of Paleovalley in Virginia Mountains A short hike north of the road leads to a small exposure of Cretaceous granite. Mafic lavas of the Pyramid sequence rest directly on the granite, whereas thick Oligocene ash-flow tuffs crop out directly to the south. These relations define the northern margin of a W-SW–trending Oligocene paleovalley. The ash-flow tuff sequence in this paleovalley is essentially identical to that found in the Nightingale Mountains (Fig. 4), indicating that these paleovalley segments were originally part of the same drainage system. Assuming a consistent W-SW trend, the paleovalley is offset ~10 km across the Pyramid Lake fault (Fig. 5, segments 1A and 1B, and Fig. 7). Directions to Stop 1-6 Return to Hwy 445, turn right, and proceed south 1.8 mi, then turn right (at the south end of the town of Sutcliffe) onto a paved road to the Dunn Fish Hatchery. Veer left around the fish hatchery onto the right fork of a dirt road, which leads into Hardscrabble Canyon. About 1.0 mi from Hwy 445, you will come to a locked gate to private property; access past this point is by permission only. We will pass through the gate and travel ~0.4 mi to Stop 1-6. Stop 1-6 (optional): Axis of Paleovalley in Virginia Mountains, Hardscrabble Canyon A short hike to the north brings us to outcrops of the 31.2 Ma tuff of Hardscrabble Canyon and 31.0 Ma tuff of Rattlesnake Canyon. The southern Virginia Mountains, including the Hardscrabble Canyon area, contain a >600-m-thick sequence of Oligocene ash-flow tuff along the axis of a major paleovalley (Faulds et al., 2003a, 2003b, 2005b). The units here are the same as those exposed at the base of the tuff sequence along the northern margin of the paleovalley in the Nightingale Mountains. The Hardscrabble Canyon area is cut by several closely-spaced, relatively minor NW-striking dextral-normal faults that are now inactive (Faulds et al., 2003a). This fault system may reflect early stage Riedel shears that predate development of more throughgoing faults (e.g., Pyramid Lake fault) and are analogous to the abundant NW-striking faults and lineaments near the northwest terminus of the Walker Lane in northeast California today. DAY 2: VIRGINIA MOUNTAINS TO DOGSKIN MOUNTAIN ACROSS WARM SPRINGS VALLEY FAULT Introduction On Day 2, we will continue tracking the Oligocene paleovalley system across the northern Walker Lane but, in this case, across

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the right-lateral Warm Springs Valley fault (Fig. 3). Unlike the Pyramid Lake fault, the Warm Springs Valley fault cuts through a bedrock saddle between the Virginia Mountains and Dogskin Mountain. This affords an opportunity to more closely analyze one of the major strike-slip faults in the northern Walker Lane. En Route Discussion As we depart Reno and travel north on Highway 445 (Fig. 3), we pass through a large west-tilted half graben known as Spanish Springs Valley. This is one of several northerly trending, west-tilted half grabens situated just north of Reno (Fig. 5). The half grabens are bounded by east-dipping normal faults that merge southward with the Sierra Nevada frontal fault system. At ~16 mi, note the change in topographic grain, as the northtrending ridges and half graben give way to northwest-trending basins and ranges within the northern Walker Lane. At ~19 mi, Hwy 445 passes through the northwest-trending Warm Springs Valley, which contains one of the major right-lateral faults within the northern Walker Lane, the Warm Springs Valley fault. Directions to Stop 2-1 From the UNR campus, head east on I-80 and continue 3.1 mi to the Pyramid Way (Hwy 445) exit in Sparks (Exit 18). Go north on Pyramid Way, which turns into the Pyramid Highway (Hwy 445) north of Sparks. At 21.8 mi, turn left onto a graded dirt road (Grass Valley Road); travel 1.5 mi to Flying Eagle Ranch. This is private property, so trenches are viewed by permission only. Stop 2-1 (optional): Trenches along Warm Springs Valley Fault (if still open) The Warm Springs Valley fault is a major 96-km-long right-lateral fault system in northwest Nevada and northeast California capable of producing earthquakes as large as ~M 7.3. Recent trenching along the fault shows evidence for five to eight earthquakes since ca. 21 ka, as evidenced by offset channels, liquefied deposits, and colluvial units shed from low scarps (dePolo et al., 2005). Directions to Stop 2-2 Exit the Flying Eagle Ranch and return to the graded dirt road. Continue north and west on this road, then turn right onto a secondary dirt road 1.0 mi north of Flying Eagle Ranch. After 0.1 mi, turn left as road “T”s out and continue 0.7 mi to west to Stop 2-2 (near intersection with Incandescent Canyon Road). Stop 2-2: View of Incandescent Canyon Stop 2-2 affords an excellent view of Incandescent Canyon in the Virginia Mountains to the north. The multicolored layers in Incandescent Canyon represent the upper part of the Oligocene ash-flow tuff section, including the 25.3 Ma Nine Hill Tuff (Deino, 1985), the 25.1 Ma tuff of Chimney Spring, and the

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24.9 Ma tuff of Painted Hills (Faulds et al., 2003b). The layercake stratigraphy in this area is deceptive, as cut-and-fill features characterize the ash-flow tuff sequences within the paleovalleys (Fig. 4). East of Incandescent Canyon, several ca. 23.5 Ma rhyolite domes intrude the section of ash-flow tuffs. Thick megabreccia deposits that crop out on several of the ridges were shed from these domes soon after their emplacement. Directions to Stop 2-3 Continue 1.3 mi west along the main dirt road, and then turn right and pass through gate. After 1.8 mi come to stop in lower part of Mine Canyon near the end of the passable road. Stop 2-3: Oligocene Ash-Flow Tuffs in Mine Canyon At this stop, we will traverse to the east ~1 km up through the section of Oligocene ash-flow tuffs, gaining ~250 m in elevation. Here, we find excellent exposures of the upper part of the Oligocene tuff section, which are cut by NW-striking normaldextral faults (Fig. 14). In ascending order, the section in this area includes the 30.1 Ma tuff of Cove Spring, an unnamed tuff, the 29.9 Ma tuff of Mine Canyon, the 29.3 Ma tuff of Dogskin Mountain, debris flow deposit dominated by porphyritic andesite clasts, the 25.3 Ma Nine Hill Tuff, the 25.1 Ma tuff of Chimney Spring, the 24.9 Ma tuff of Painted Hills (upper and lower units), a megabreccia deposit, and a younger tuff that may correlate with the 23.7 Ma tuff of Perry Canyon (Fig. 4; Faulds et al., 2003b). The megabreccia consists of clasts and blocks of older tuffs and porphyritic rhyolite domes up to 10 m in diameter in a coarse, granular matrix. This megabreccia formed as rock avalanches induced by catastrophic failure of steeply dipping strata around the flanks of the intrusive rhyolite domes. It is lithologically similar to intracaldera or tectonic megabreccias but does not indicate

proximity to a caldera or coeval tectonism. The tuffs in this area exceed 600 m in thickness and clearly lie within the axial part of a W-SW–trending paleovalley. The granitic massif of Dogskin Mountain (with no exposures of tuff) lies on trend with this paleovalley segment directly W-SW across the Warm Springs Valley fault (Fig. 5, segment 1B). The westward continuation of the paleovalley is found in the northwest part of Dogskin Mountain, which indicates ~10 km of dextral offset across the Warm Springs Valley fault. Just to the north, the Oligocene tuffs are capped by a thick section of middle Miocene mafic lavas within the Pyramid sequence. A northerly striking dike swarm and multiple cinder cones suggest that a relatively large shield volcano occupied the central part of the Virginia Mountains (Faulds and Henry, 2002). Much of the Pyramid sequence in this area was probably erupted from that volcano. Directions to Stop 2-4 From Stop 2-3, travel 1.8 mi back down the road, through the gate, then turn right and travel 0.2 mi. Turn left onto a small spur road toward a low ridge. We will walk up the old road to a small cut that exposes late Miocene–Pliocene sedimentary rocks. Stop 2-4: Tephra-Bearing Linear Ridge Along Warm Springs Valley Fault System This ridge, which rises ~140 m above the surrounding alluvial fans, is the southernmost of a series of linear ridges along the Warm Springs Valley fault system (Fig. 15; Faulds et al., 2003b; Henry et al., 2006). A prominent strand of the fault lies along the northeast side of the ridge and can be seen to continue to the northwest. Geologic relationships indicate that a parallel strand lies southwest of the ridge. The ridge is cored by late Miocene–Pliocene sedimentary rocks, which consist here of coarse sandstone

Figure 14. View looking east at Oligocene ash-flow tuffs in Mine Canyon, which include the tuff of Dogskin Mountain (Td), Nine Hill Tuff (Tnh), tuff of Chimney Spring (Tcs), tuff of Painted Hills (Tph; two units), and a capping megabreccia deposit (Tbt). The tuffs are cut by a NW-striking, down-to-thesouthwest normal-dextral fault. Stop 2-3 traverses up through this section.

Transect across the northern Walker Lane and pebbly conglomerate containing clasts of Cretaceous quartz monzodiorite similar to that making up Dogskin Mountain to the west. Clasts are almost entirely less than a few centimeters in diameter. The small excavation provides the only good exposure of the deposits and shows two dacitic tephra containing phenocrysts of plagioclase and biotite. A single crystal 40Ar/39Ar date on plagioclase from the upper tephra is 3.58 ± 0.45 Ma (Henry et al., 2006). This and similar tephra in the sedimentary sequence throughout the region were probably erupted from volcanoes of the ancestral Cascade arc not far to the west in California. Several features suggest that a major episode of normal faulting postdates ca. 3.6 Ma (Henry et al., 2006). First, the Pliocene sedimentary rocks at this stop and Mio-Pliocene sedimentary rocks in much of the region are tilted about the same amount and in the same direction as both the 15–14 Ma Pyramid sequence and 31–23 Ma ash-flow tuffs. Second, the relatively fine grain size of the sedimentary rocks implies that the present, high Dogskin Mountain did not exist when they were deposited (although climatic influences cannot be ruled out; e.g., Molnar, 2004). A low “ancestral Dogskin” ridge or more distal outcrop of quartz monzodiorite must have existed to provide detritus, but modern alluvial fans on the flanks of Dogskin Mountain are coarse down to the fault zone and axial drainage, with clasts commonly to 1 m in diameter. Although this interpretation is somewhat controversial even among your trip leaders, it suggests that most uplift of Dogskin Mountain occurred after 3.6 Ma, probably during the regional episode of east-west extension at ca. 3 Ma (Henry and Perkins, 2001). Strike-slip faulting then began after the episode of extension, after ca. 3 Ma.

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Directions to Stop 2-6 Continue on graded road for another 5.5 mi, then turn left onto a secondary dirt road, and travel 0.5 mi to Stop 2-6 (just before the road becomes nearly impassable at “+” in Fig. 16). Stop 2-6: McKissick Canyon—Complex Paleovalley Cut and Fill We will make a short (~1 km) traverse to examine complex thickening and thinning of ash-flow tuffs in the paleovalley. This location ~1.5 km northwest of the paleovalley wall shows complex cut and fill between the tuff of Dogskin Mountain, tuff of Campbell Creek, sedimentary breccia, Nine Hill Tuff, tuff of Chimney Spring, and lower tuff of Painted Hills (Fig. 16; Henry et al., 2004b). First, an intra-tuff paleovalley was eroded at least 100 m deep into the 29.3 Ma tuff of Dogskin Mountain. This valley was mostly filled by 28.8 Ma tuff of Campbell Creek, which was then substantially eroded. The resulting valley was partly

Directions to Stop 2-5 Return to the main dirt road, turn left, and travel west 1.3 mi. Turn right onto a graded county road, which follows the Warm Springs Valley fault system, and travel 1.0 mi to Stop 2-5, a low roadcut on the northeast side of the road. Stop 2-5: Vertical Beds along Warm Springs Valley Fault System The roadcut exposes late Miocene–Pliocene sedimentary rocks, which here strike northwest and are vertical. The beds consist of pebbly conglomerate composed of clasts of ash-flow tuff and lesser Cretaceous quartz monzodiorite. The change in clast type suggests these rocks were deposited farther northwest, near where ash-flow tuff is exposed in Dogskin Mountain. The fine grain size again suggests an “ancestral” low Dogskin Mountain that has been uplifted since deposition. If the sedimentary rocks were deposited near the occurrence of tuff in Dogskin Mountain, they have been translated 4–5 km southeastward along a strand of the fault system southwest of here. The prominent strand seen at stop 4 lies ~250 m to the northeast. The steep tilt of beds here is consistent with their occurrence between the major strands.

Figure 15. Oblique aerial view looking northwest along the Warm Springs Valley fault and linear ridges. Stop 2-4 is in ridge in foreground, which contains the 3.58 Ma tephra at its northwest end. Prominent linear trend along northeast side of ridges is northeastern splay of the Warm Springs Valley fault, which does not cut youngest alluvial fan. Another splay runs southwest of the ridges. Apparent layers on southwest side of the ridge are Lake Lahontan shorelines. The Warm Springs Valley fault continues northwestward into highlands in the distance, the one area where northern Walker Lane faults are expressed in bedrock.

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filled by sedimentary breccia. This was also eroded, and Nine Hill Tuff filled that valley. Each unit is thick in the middle of each intra-tuff channel but thins and pinches out short distances laterally. For example, the Nine Hill Tuff is ~50 m thick in the canyon at the location marked “+” on Figure 16, but thins to ~5 m ~250 m to the southeast, and pinches out another 200 m farther southeast from there and also 200 m northwest of the canyon. Much of the thinning is depositional; that is, the base of the Nine Hill climbs over an irregular upper surface of tuff of Campbell Creek and sedimentary breccia. However, the lack of an upper, nonwelded zone of the Nine Hill Tuff shows that it was also eroded and that its initial thickness was considerably greater than that preserved. Abrupt contacts between different units that strike into each other could easily be misinterpreted as faults. Two units are notable. (1) The tuff of Campbell Creek erupted from a caldera in the Desatoya Mountains, ~190 km to the east-southeast. Outflow Campbell Creek is present in the Clan Alpine Mountains ~180 km east of here. However, we have not found it between those eastern areas and Dogskin Mountain, despite detailed mapping in the Virginia Mountains and substantial reconnaissance in the Lake and Nightingale Ranges. Yet from here west, the tuff of Campbell Creek is prominent, and we have found it as far west as ~20 km east of Nevada City, California, in the Sierra Nevada foothills ~100 km southwest of here. These relationships mean that it was deposited throughout that length but almost entirely eroded over most of its distribution. (2) A sedimentary breccia is discontinuously exposed for ~7 km along what must be the remnants of an intra-tuff channel along

the southeast side of the main Dogskin paleovalley. Clasts here are of tuff of Dogskin Mountain up to ~2 m in diameter and of probable Mesozoic porphyritic andesite. Where better exposed to the southwest, the gravel is matrix-supported with angular to subrounded clasts up to 12 m in diameter. These characteristics suggest deposition as debris flows. Clasts almost universally have finely brecciated but recemented margins covering intact cores, similar to clasts in rock-avalanche deposits. Directions to Stop 2-7 Return to the county road, turn left, and travel 0.25 mi to Stop 2-7 along the south side of the road where side road leads southwest. Stop 2-7: Rhyolite Dome-Related Megabreccia This stop shows megabreccia similar to that observed at Stop 2-3. Megabreccia closest to the road consists of clasts of porphyritic rhyolite-dacite, either of the domes or of ash-flow tuff related to them. Other megabreccia is dominated by clasts of tuff of Chimney Spring up to several meters in diameter. Megaclasts of Nine Hill Tuff up to 200 m long lie less than 1 km to the north. All known outcrops of the domes are northeast of the Warm Springs Valley fault system, mostly near Incandescent Canyon. Presence of megabreccia here indicates this location was probably near those domes before displacement along the Warm Springs Valley fault. Directions to Stop 2-8 Continue along graded county road for another 1.1 mi to prominent saddle, which is Stop 2-8. Hike uphill to the north of the road. Stop 2-8: Panoramic Views of the Warm Springs Valley Fault System

Figure 16. Complex thickness variations of ash-flow tuffs resulting from sequential deposition and erosion in intra-tuff channels in the Dogskin Mountain paleovalley (Henry et al., 2004b). Tbt—rhyolite dome breccia; Tphl—tuff of Painted Hills; Tcs—tuff of Chimney Spring; Tnh—Nine Hill Tuff; Tcn—sedimentary breccia; Tcc—tuff of Campbell Creek; Tdm—tuff of Dogskin Mountain. Numbers next to unit symbols are 40Ar/39Ar ages in Ma.

This stop provides excellent panoramic views of several features of the Warm Springs Valley fault system. Outcrop is of Nine Hill Tuff, which here forms a NW-trending syncline between branches of the fault system (Henry et al., 2004b). In part, the syncline is extensional and predated strike-slip faulting. Extension tilted rocks northeast along the northeast flank of Dogskin Mountain and west along the west flank of the Virginia Mountains. However, strike-slip faulting has greatly accentuated the dips to the point that the east flank of the syncline is locally overturned. The view to the southeast shows the series of linear ridges along the fault system and its parallelism with the Dogskin Mountain range-front fault. The latter fault makes a prominent topographic scarp but shows no evidence of Quaternary movement. The Warm Springs Valley fault lies 1–2 km northeast of the Dogskin Mountain fault and must either cut or reactivate it at depths of no more than a few km. Right-lateral displacement of the paleovalley from the Mine Canyon area of the Virginia Mountains to Dogskin Mountain is apparent.

Transect across the northern Walker Lane The view to the northwest shows the fault system continuing in bedrock for ~20 km, the only part of a northern Walker Lane fault not in a basin. This bedrock part of the fault system strikes ~N55°W and changes to N40°W southward across Dry Valley Creek just below us to the north. The change in strike coincides with an ~3.5-km-long, east-striking fault or fold zone that crosses the fault system (Fig. 5 and Henry et al., 2004b). This zone is along a particularly linear, east-trending part of Dry Valley Creek that is distinctly aberrant to regional drainage patterns and extends another 4 km beyond the western end of the cross structure. Along most of the western half of the cross structure, volcanic rocks strike ~E on both the north and south sides, in contrast to their regional north to northwest strikes. A W-NW–trending syncline is slightly oblique to the cross structure on its south side. Along the eastern part of the cross structure, volcanic rocks on the north side maintain their N-NW strike but are truncated against gently south-dipping rocks on the south side. Expression of the Warm Springs Valley fault system in bedrock and the apparently contractional structures across Dry Valley Creek may reflect the probable transpressional, N55°W segment of the Warm Springs Valley fault system, which is oblique to the ~N40–45°W direction of relative transport between the Sierra Nevada and western Great Basin (Hammond et al., 2004). DAY 3: SEVEN LAKES MOUNTAIN TO DIAMOND MOUNTAINS ACROSS HONEY LAKE FAULT On Day 3, we will continue tracking the Oligocene paleovalley system westward to the Nevada-California border area near Seven Lakes Mountain and then to the Diamond Mountains in northeast California (Fig. 3). We will evaluate constraints on offset across the Honey Lake fault zone, the third in the series of left-stepping right-lateral faults in the northern Walker Lane. Directions to Stop 3-1 From the UNR campus, head north on Virginia Street ~4.2 mi and merge with Hwy 395. Continue north on U.S. 395 through Long Valley (Fig. 3), which contains Mio-Pliocene sedimentary rocks tilted westward into normal faults along the west side. After 28.7 mi, turn right onto a dirt road. After 1.4 mi, veer left just after gate, then continue 0.1 mi and stop at the top of a ridge. Stop 3-1: Paleovalley Margin at North End of Peterson Mountain At this stop, we view the southern margin of a tuff-filled paleovalley exposed at the northern end of Peterson Mountain. We correlate this paleovalley margin with that exposed at Dogskin Mountain (near Stop 2-6). The tuff of Sutcliffe was deposited against granodiorite of Peterson Mountain in this area. The contact roughly follows the gully and road that lead eastward down to Red Rock Valley. About 100 m north of the road, dense, devitrified tuff of Sutcliffe has an attitude of ~N10°E, 14°W, similar to that farther north. Next to the road near the top of the

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gully, the tuff of Sutcliffe includes vitrophyre and has an attitude of N30°E, 27°W. The presence of vitrophyre and the slightly different attitude probably result from cooling and compaction against a moderately steep wall cut in granodiorite. Near the bottom of the gully, the tuff of Sutcliffe forms what appears to be a north-dipping slope against granodiorite on the south side of the gully. The tuff does not crop out, but pieces of both vitrophyre and devitrified rock are present. These relationships indicate the contact is depositional and not a fault. Thus, this location is the south margin of the paleovalley. It is approximately on trend with the south margin at Dogskin Mountain, which indicates only a small amount, probably a few km, of cumulative dextral slip across the southern part of the Honey Lake fault zone. Directions to Stop 3-2 From Stop 1, continue east on the dirt road for 0.4 mi, then turn left onto paved Red Rocks Road. Continue west on the paved road for 2.2 mi and turn right into pullout. Stop 3-2: Paleovalley and South Limb of Seven Lakes Mountain Anticline We will hike ~1 km northward to view (1) the ash-flow tuff section somewhat farther into the paleovalley than at stop 1, and (2) steep dips of the tuffs and overlying sedimentary rocks on the south limb of the Seven Lakes Mountain anticline (Fig. 5) (Henry et al., 2005). Exposed rocks are, from oldest to youngest: coarse, euhedral biotite-hornblende granodiorite, probable tuff of Rattlesnake Canyon, tuff of Sutcliffe, conglomerate and sandstone, probable tuff of Cove Spring, lower and upper facies of Nine Hill Tuff, and late Miocene–Pliocene, basin-filling sedimentary rocks. Below the tuff of Rattlesnake Canyon, a few rounded cobbles and pebbles of granitic rock and tuff of Axehandle Canyon (?) suggest the presence of a post-Axehandle conglomerate. Nine Hill Tuff is exposed only at the west end of this ridge, where it fills an intra-tuff canyon cut entirely through the older tuffs. South of the Red Rocks Road and to the east of our hike north of the road, sedimentary rocks and ash-flow tuffs dip 15° to 25° westward. The rocks turn abruptly to W-NW strikes and steep (55° to 75°) south dips along the south flank of Seven Lakes Mountain and the south limb of the anticline. All rocks are similarly tilted, which suggests that most deformation postdates deposition of the late Miocene–Pliocene sedimentary rocks. However, the sedimentary deposits include coarse conglomerate containing rounded boulders of granitic rock to ~1 m, which suggests the presence of steep topography, probably a nearby fault scarp, at the time of deposition. The fault scarp probably formed during the episode of low-magnitude extension at ca. 12 Ma (Henry and Perkins, 2001; Henry et al., 2006). Eastward along this south limb, probable tuff of Cove Spring is in contact with the younger tuff of Dogskin Mountain, both steeply south-dipping, along a shallowly north-dipping fault. Tuff of Cove Spring is intensely brecciated along

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the fault. In its present attitude, this fault is a thrust, which is consistent with folding of Seven Lakes Mountain. However, the fault could be normal, south side down, if it formed when the tuffs were approximately flat-lying. Directions to Stop 3-3 Turn right onto Red Rocks Road from pullout and proceed 0.2 mi to Hwy 395. Turn right onto the highway, travel north 0.9 mi, and turn right onto a dirt road to communication towers. After 1.1 mi, stop to view outcrops along and near the road. Stop 3-3: North Limb of Anticline, Northwest Side of Seven Lakes Mountain At this stop, we will view north-dipping ash-flow tuffs on the north limb of the Seven Lakes Mountain anticline. The tuff of Sutcliffe strikes E-NE and dips ~60° northward at this location. The tuff rests on granodiorite but is distinguished from the petrographically similar but older tuff of Axehandle Canyon because pebbles of tuff of Hardscrabble Canyon appear to underlie the tuff. Probable Te, tuff of Campbell Creek, and late Miocene–Pliocene sedimentary rocks overlie the tuff below us. The low foothills just to the north are underlain by the sedimentary rocks, which are folded into a syncline and another anticline. Still farther north, Dry Valley is a large syncline, with a thin veneer of Quaternary deposits over the sedimentary rocks. All folds trend ~W to W-NW. Directions to Stop 3-4 Continue on the communications road to the top of mountain 1.1 mi past Stop 3-3. Stop 3-4: Axis of Paleovalley and Crest of Anticline, Seven Lakes Mountain This stop affords a good view of the Fort Sage Mountains and Honey Lake basin. We will take a short traverse around the top of the ridge to look at gently dipping tuffs that lie in the hinge zone of the anticline. Exposed rocks are the coarse, euhedral biotite-hornblende granodiorite and probable tuffs of Axehandle Canyon, Hardscrabble Canyon, and Rattlesnake Canyon. All three tuffs rest on granodiorite next to each other and do not make an obvious stratigraphic succession. Identification is based on phenocryst assemblages. Additionally, the tuff of Rattlesnake Canyon contains a lens of vitrophyre adjacent to its contact with the tuff of Axehandle Canyon, which suggests the former overlies the latter. The presence of the oldest tuffs here suggests that we are at or near the axis of the paleovalley. However, the tuff sequence is covered along much of the north flank of Seven Lakes Mountain and in Dry Valley to the north, so the exact location of the axis is uncertain. The top of the ridge affords a good view northward to Dry Valley, the Fort Sage Mountains, and along the approximate trace

of the Honey Lake fault system, and west- and southwestward to the Sierra Nevada and Long Valley, a west-tilted half graben that appears to splay from the Honey Lake fault. The Fort Sage Mountains are a gently (~6°) west-tilted fault block that contains a west-trending tributary paleovalley, with an ~200 m thick sequence of ash-flow tuffs, including the tuff of Axehandle Canyon, tuff of Campbell Creek, Nine Hill Tuff, and tuff of Chimney Spring (Fig. 5, segment 3A; Hinz, 2004). Although not well expressed in the valley, the Honey Lake fault strikes southeastward toward the middle of Seven Lakes Mountain and must make a substantial left step around the mountain. Contraction in this left step has generated the Seven Lakes Mountain anticline and Dry Valley syncline. Displacement on the Honey Lake fault system south of Seven Lakes Mountains is only ~3–5 km, substantially less than the ~10 km on the northern, main part of the Honey Lake fault to the northwest. Part of the difference could be taken up by extension along the north-striking normal faults that bound the west side of Long Valley. We are currently mapping the Seven Lakes Mountain Quadrangle to understand the geometry, displacement, and evolution of the fault system, particularly how or whether the two parts of the fault system are linked. Directions to Stop 3-5 From Stop 3-4, return to U.S. 395 and turn right. Continue NW on Hwy 395 for 24.7 mi and turn left at Milford Grade turnoff. After 1.5 mi, turn left onto Milford Grade and continue upgrade and through a pass at 5.4 mi, then turn left onto Black Mountain Road and stop after traveling 1.1 mi. Stop 3-5: Oligocene Paleovalley at Black Mountain in Diamond Mountains, California At this stop, the tuffs of Axehandle Canyon and Rattlesnake Canyon in the lower part of the Oligocene section are well exposed just a short distance north of the road. Correlation of the tuffs here was based initially on stratigraphic position and phenocryst assemblage and has been confirmed by both 40Ar/39Ar dates and paleomagnetism. The tuff of Axehandle Canyon is only ~10 m thick here and crops out over a small area. However, it can be traced discontinuously ~6 km eastward to the east flank of the Diamond Mountains above Honey Lake (Hinz, 2004). The track appears to mark the bottom of this paleovalley. The tuff of Rattlesnake Canyon (Fig. 17) is much more widely distributed and up to ~125 m thick (Hinz, 2004). It shows moderately good columnar joints and a thick basal vitrophyre here. About 2 km to the northeast, the tuff makes two distinct but petrographically similar ledges, which suggest a compound cooling unit consistent with its location here far from a possible caldera source. The near horizontal average of strikes and dips of the tuff sequence around Black Mountain indicate negligible tilting of the structural block (Hinz, 2004). These dips are typically <15° but are locally as steep as 35°. Because the Diamond Mountains are at most very gently tilted, these dips indicate compaction against existing topography and do not record paleohorizontal.

Transect across the northern Walker Lane The thickness and sequence of tuffs in the Black Mountain area is very similar to that in the Fort Sage Mountains, suggesting that the paleovalley segments in each were once continuous. On the basis of their location and thinner cumulative sequence of tuff, we interpret the paleovalley segments in the Fort Sage and Diamond Mountains as a tributary to the main-stem paleovalley exposed in the Nightingale Mountains, Virginia Mountains, Dogskin Mountain, and Seven Lakes Mountain (Fig. 5). Compared to paleovalley segments farther east, the areal extent of exposures is limited in the Fort Sage and Diamond Mountains. The trend of the paleovalleys is therefore difficult to constrain, which in turn makes estimating offset on the Honey Lake fault more equivocal. The limited exposures and anisotropy of magnetic susceptibility data suggest SW- to W-trending paleovalleys in both the Fort Sage and Diamond Mountains, which implies 9–17 km of displacement along the Honey Lake fault zone (Hinz, 2004). Considering the relatively small amount of offset on the southeastern part of the Honey Lake fault zone and the more westerly trends of other paleovalley segments, we conclude that the lower estimate of offset is more likely. Directions to Stop 3-6 Continue up Black Mountain Road several miles to an exposed northern slope with a good view of Honey Lake basin. Stop 3-6: North Side of Black Mountain, View of Honey Lake Basin This final stop provides excellent views of the Honey Lake basin, Honey Lake fault zone, and Diamond Mountains rangefront fault. The Honey Lake fault system roughly parallels the Diamond Mountains fault for ~35 km but is 1–5 km out into the hanging wall. Honey Lake basin is a major basin, at least 3 km deep based on gravity data (Ceron, 1991), which requires substantial vertical displacement along the Diamond Mountains fault. The shape of the basin, the position of the Honey Lake fault system well into the basin, and the amount of vertical displacement relative to probable strike-slip displacement suggest that the basin resulted primarily from extension and vertical displacement along the Diamond Mountains fault, not from strike-slip displacement along the Honey Lake fault system. However, the Diamond Mountains fault has no Quaternary scarps (Grose et al., 1990; U.S. Geological Survey, 2004), which suggests that it is inactive. In contrast, the Honey Lake fault system is clearly active. Grose (2000) and Wills and Borchardt (1993) identified discontinuous Holocene scarps along a 30-km-long segment of the Honey Lake fault system stretching southeastward from Honey Lake and adjacent to the Diamond Mountains fault. Wills and Borchardt (1993) estimated a Holocene slip rate of between 1.1 and 2.6 mm/yr. Highly deformed, late Pliocene sedimentary rocks are spectacularly exposed along the Honey Lake fault system along the northwest side of the peninsula extending into Honey Lake, as can be seen to the north (Grose et al., 1990; Wills and Borchardt,

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1993; Adams et al., 2001; Mass et al., 2003; Park et al., 2004). The sedimentary rocks are lacustrine sandstones, siltstones, and diatomite with numerous tephra. These deposits crop out between two major splays of the Honey Lake fault; the northeastern splay has a 2–3-m-high scarp, down to the northeast, in Holocene Lake Lahontan deposits (Grose, 2000; Adams et al., 2001). The sedimentary rocks are complexly deformed with as many as four high-angle faults, presumably splays of the Honey Lake fault, recumbent folds, and thrust faults in a local zone of transpression (Adams et al., 2001; Mass et al., 2003; Park et al., 2004). The sedimentary rocks show only minor evidence for deformation during deposition. For example, Mass et al. (2003) interpreted soft-sediment slumping low in the section as possible coseismic deformation. However, angular unconformities, as would be expected if the sedimentary rocks had been deformed along the Honey Lake fault during deposition and as now is present between the Pliocene strata and Holocene Lake Lahontan deposits, are absent. We attempted to identify and date the stratigraphically highest and lowest tephra, although the complex deformation allows considerable uncertainty in interpreting the stratigraphic section. Biotite 40Ar/39Ar ages of the two tephra are 2.95 ± 0.13 Ma and 3.71 ± 0.10 Ma (plateau) and 2.91 ± 0.05 Ma and 3.73 ± 0.03 Ma (isochron) (Henry et al., 2006). Our initial interpretation of the field and 40Ar/39Ar data is that the sedimentary rocks accumulated in Honey Lake basin entirely before the onset of strike-slip faulting. Basin formation and deposition certainly started long before 3.7 Ma, with a probable maximum age of 16 Ma (Wagner et al., 2000). The softsediment deformation identified by Mass et al. (2003) may have resulted from normal faulting on the Diamond Mountains fault but is far less intense than the folding and thrust-faulting presumably generated by the subsequent strike-slip faulting. Therefore,

Figure 17. Oligocene ash-flow tuffs in the Black Mountain area of the Diamond Mountains. Kqmd—Cretaceous quartz monzodiorite; Trc—31.0 Ma tuff of Rattlesnake Canyon; Ts—30.4 Ma tuff of Sutcliffe; Tcc—two cooling units of the 28.8 Ma tuff of Campbell Creek (Hinz, 2004).

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strike-slip faulting along the Honey Lake fault probably began after 2.9 Ma. An alternative that allows earlier strike-slip faulting would require initial displacement along the Honey Lake fault on strands that were more distal from the exposed deposits or that accommodated pure strike-slip displacement with negligible contractional deformation. Exposures of the Warm Springs Valley fault in bedrock (Stop 2-8) demonstrate that strike-slip faults of the northern Walker Lane can form zones as much as 1.5 km wide (Henry et al., 2004b). The exposed Honey Lake fault zone is ~1 km wide, and even allowing for an early fault to be another 1.5 km away may not have been sufficient to prevent the kind of deformation shown by the sedimentary rocks. Continuing investigation of these deposits and their deformation (Mass et al., 2003; Park et al., 2004) should test these alternatives. SUMMARY East of the San Andreas fault within the western Great Basin, a system of dextral strike-slip faults accommodates a significant fraction of the North American–Pacific plate motion. The northern Walker Lane occupies the northern terminus of this fault system and is one of the youngest and least developed parts of the North American–Pacific transform plate boundary. Accordingly, the northern Walker Lane affords an opportunity to analyze the incipient development of a major strike-slip fault system. In northwest Nevada, the northern Walker Lane consists of an ~50-km-wide belt of overlapping, curiously left-stepping dextral faults, whereas a much broader zone of disconnected, widely-spaced northwest-striking faults characterizes northeast California. The left steps accommodate little shortening and are not typical restraining bends. The left-stepping dextral faults may represent Riedel shears developing above a nascent lithosphericscale transform fault. The widely-spaced faults in northeast California may also be Riedel shears but at an earlier developmental stage with no through-going strike-slip fault at depth. Strands of the northern Walker Lane terminate in arrays of northerly striking normal faults in the northwestern Great Basin and along the eastern front of the Sierra Nevada, suggesting that dextral shear is transferred to ~NW-SE extension in the Great Basin (Fig. 11). On this trip, we have followed an Oligocene paleovalley system across nearly the entire northern Walker Lane (Fig. 5). Inferred offsets of the paleovalley system indicate ~20–30 km of cumulative dextral displacement, as measured orthogonal to the northern Walker Lane. Strike-slip faulting began between 3 and 9 Ma, indicating a long-term slip rate of ~2–10 mm/yr, which is compatible with GPS geodetic observations of the current strain field. ACKNOWLEDGMENTS This work was supported by National Science Foundation grant EAR-0124869 and the STATEMAP Program of the U.S. Geological Survey. We thank John Bell, Woody Brooks, Patricia Cashman, Craig dePolo, Larry Garside, Mark Hudson, John

Oldow, George Saucedo, and Dave Wagner for fruitful discussions about the northern Walker Lane. We also thank Kris Pizarro for preparation of the figures. Constructive reviews by Susanne Janecke and Joel Pederson significantly improved this field trip guide. REFERENCES CITED Adams, K.D., Wesnousky, S.G., and Bills, B.G., 1999, Isostatic rebound, active faulting, and potential geomorphic effects in the Lake Lahontan basin, Nevada and California: Geological Society of America Bulletin, v. 111, p. 1739–1756, doi: 10.1130/0016-7606(1999)111<1739: IRAFAP>2.3.CO;2. Adams, K., Briggs, R., Bull, W., Brune, J., Granger, D., Ramelli, A., Riebe, C., Sawyer, T., Wakabayashi, J., and Wills, C., 2001, Northern Walker Lane and northeast Sierra Nevada: Friends of the Pleistocene Pacific Cell Field Trip, 67 p. An, L.-J., and Sammis, C.G., 1996, Development of strike-slip faults: Shear experiments in granular materials and clay using a new technique: Journal of Structural Geology, v. 18, p. 1061–1078, doi: 10.1016/0191-8141(96)00012-0. Armijo, R., Fleri, F., King, G., and Meyer, B., 2004, Linear elastic fracture mechanics explains the past and present evolution of the Aegean: Earth and Planetary Science Letters, v. 217, p. 85–95, doi: 10.1016/S0012821X(03)00590-9. Atwater, T., and Stock, J., 1998, Pacific-North America plate tectonics of the Neogene southwestern United States: An update: International Geology Review, v. 40, p. 375–402. Baljinnyam, I., Bayasgalan, A., Borisov, B.A., Cisternas, A., Dem Yanovich, M.G., Ganbaatar, L., Kochetkov, V.M., Kurushin, R.A., Molnar, P., Philip, H., and Vashchilov, Y.Y., 1993, Ruptures of major earthquakes and active deformation in Mongolia and its surroundings: Geological Society of America Memoir 181, 62 p. Bayasgalan, A., Jackson, J., Ritz, J.-F., and Carretier, S., 1999, Field examples of strike-slip fault terminations in Mongolia and their tectonic significance: Tectonics, v. 18, p. 394–411, doi: 10.1029/1999TC900007. Bell, J.W., 1984, Quaternary fault map of Nevada, Reno sheet: Nevada Bureau of Mines and Geology Map 79, scale 1:250,000. Bell, J.W., and House, P.K., 2005, Pattern and role of late Quaternary faulting along the Pyramid Lake fault zone, northern Walker Lane belt, based on deformation of Lake Lahontan chronostratigraphy: Seismological Society of America Research Letters, v. 76, no. 2, p. 245. Bennett, R.A., Wernicke, B.P., Niemi, N.A., Friedrich, A.M., and Davis, J.L., 2003, Contemporary strain rates in the northern Basin and Range province from GPS data: Tectonics, v. 22, p. 3-1–3-31. Benson, L., 1994, Carbonate deposition, Pyramid Lake subbasin, Nevada: Sequence of formation and elevational distribution of carbonate deposits (tufas): Palaeoecology, v. 109, p. 55–87. Best, M.G., Christiansen, E.H., Deino, A.L., Grommé, C.S., McKee, E.H., and Noble, D.C., 1989, Eocene through Miocene volcanism in the Great Basin of the western United States: New Mexico Bureau of Mines and Mineral Resources Memoir 47, p. 91–133. Bonham, H.F., and Papke, K.G., 1969, Geology and mineral deposits of Washoe and Storey Counties, Nevada: Nevada Bureau of Mines and Geology Bulletin 70, 139 p. Briggs, R.W., and Wesnousky, S.G., 2004, Late Pleistocene fault slip rate, earthquake recurrence, and recency of slip along the Pyramid Lake fault zone, northern Walker Lane, United States: Journal of Geophysical Research, v. 109, p. B08402, doi: 10.1029/2003JB002717. Brooks, E.R., Wood, M.M., Boehme, D.R., Potter, K.L., and Marcus, B.I., 2003, Geologic map of the Haskell Peak area, Sierra County, California: California Geological Survey Map Sheet 55. Cashman, P.H., and Fontaine, S.A., 2000, Strain partitioning in the northern Walker Lane, western Nevada and northeastern California: Tectonophysics, v. 326, p. 111–130, doi: 10.1016/S0040-1951(00)00149-9. Ceron, J.F., 1991, Gravity modeling of the Honey Lake basin [M.S. thesis]: Colorado School of Mines, 145 p. Christiansen, R.L., and Yeats, R.S., 1992, Post Laramide geology of the U.S. Cordilleran region, in Burchfiel, B.C., Lipman, P.W. and Zoback, M.L., eds., The Cordilleran orogen: Conterminous U.S.: Boulder, Colorado,

Transect across the northern Walker Lane Geological Society of America, The Geology of North America, v. G-3, p. 261–406. Colie, E.M., Roeske, S., and McClain, J., 2002, Strike-slip motion on the northern continuation of the late Cenozoic to Quaternary Honey Lake fault zone, Eagle Lake, northeast California: Geological Society of America Abstracts with Programs, v. 34, no. 5, p. 108. Deino, A.L., 1985, Stratigraphy, chemistry, K-Ar dating, and paleomagnetism of the Nine Hill Tuff, California-Nevada; Miocene/Oligocene ash-flow tuffs of Seven Lakes Mountain, California-Nevada; improved calibration methods and error estimates for potassium-40-argon-40 dating of young rocks [Ph.D. dissertation]: Berkeley, University of California, 457 p. dePolo, C.M., Anderson, J.G., dePolo, D.M., and Price, J.G., 1997, Earthquake occurrence in the Reno-Carson City urban corridor: Seismological Research Letters, v. 68, p. 386–397. dePolo, C.M., Sawyer, T.L., Ramelli, A.R., and Berger, G.W., 2005, Recent paleoearthquake history of the southern Warm Springs Valley fault system, northwest Nevada: Seismological Society of America Research Letters, v. 76, no. 2, p. 240. Dilek, Y., and Moores, E.M., 1999, A Tibetan model for the early Tertiary western United States: Geological Society [London] Journal, v. 156, p. 929–941. Dilles, J.H., and Gans, P.B., 1995, The chronology of Cenozoic volcanism and deformation in the Yerington area, western Basin and Range and Walker Lane: Geological Society of America Bulletin, v. 107, p. 474–486, doi: 10.1130/0016-7606(1995)107<0474:TCOCVA>2.3.CO;2. Dixon, T.H., Miller, M., Farina, F., Wang, H., and Johnson, D., 2000, Presentday motion of the Sierra Nevada block and some tectonic implications for the Basin and Range province: North American Cordillera: Tectonics, v. 19, p. 1–24, doi: 10.1029/1998TC001088. Dokka, R.K., and Travis, C.J., 1990, Late Cenozoic strike-slip faulting in the Mojave Desert, California: Tectonics, v. 9, p. 311–340. Ekren, E.B., and Byers, F.M., Jr., 1984, The Gabbs Valley Range—a wellexposed segment of the Walker Lane in west-central Nevada, in Lintz, J., Jr., ed., Western Geologic Excursions: Reno, Nevada, Geological Society of America, Guidebook 4, p. 204–215. Faulds, J.E., and Henry, C.D., 2002, Tertiary stratigraphy and structure of the Virginia Mountains, western Nevada: Implications for development of the northern Walker Lane: Geological Society of America Abstracts with Programs, v. 34, no. 5, p. 84. Faulds, J.E., dePolo, C.M., and Henry, C.D., 2003a, Preliminary geologic map of the Sutcliffe Quadrangle, Washoe County, Nevada: Nevada Bureau of Mines and Geology Open-File Report 03-17, scale 1:24,000. Faulds, J.E., Henry, C.D., and dePolo, C.M., 2003b, Preliminary geologic map of the Tule Peak Quadrangle, Washoe County, Nevada: Nevada Bureau of Mines and Geology Open-File Report 03-10, scale 1:24,000. Faulds, J.E., Henry, C.D., and Hinz, N.H., 2003c, Kinematics and cumulative displacement across the northern Walker Lane, an incipient transform fault, northwest Nevada and northeast California: Geological Society of America Abstracts with Programs, v. 35, no. 6, p. 305. Faulds, J.E., Coolbaugh, M., Blewitt, G., and Henry, C.D., 2004a, Why is Nevada in hot water? Structural controls and tectonic model of geothermal systems in the northwestern Great Basin: Geothermal Resources Council Transactions, v. 28, p. 649–654. Faulds, J.E., Henry, C.D., Hinz, N.H., Delwiche, B., and Cashman, P.H., 2004b, Kinematic implications of new paleomagnetic data from the northern Walker Lane, western Nevada: Counterintuitive anticlockwise verticalaxis rotation in an incipient dextral shear zone: Eos (Transactions, American Geophysical Union), v. 85, no. 47, Abstract GP42A-08. Faulds, J.E., Henry, C.D., Coolbaugh, M.F., Garside, L.J., and Castor, S.B., 2005a, Late Cenozoic strain field and tectonic setting of the northwestern Great Basin, western USA: Implications for geothermal activity and mineralization: Geological Society of Nevada Symposium (in press). Faulds, J.E., Henry, C.D., and Hinz, N.H., 2005b, Kinematics of the northern Walker Lane: An incipient transform fault along the Pacific–North American plate boundary: Geology, v. 33, p. 505–508, doi: 10.1130/G21274.1. Garside, L.J., Castor, S.B., Henry, C.D., and Faulds, J.E., 2000, Structure, volcanic stratigraphy, and ore deposits of the Pah Rah Range, Washoe County, Nevada: Geological Society of Nevada Symposium 2000 Field Trip Guidebook no. 2, 132 p. Garside, L.J., Henry, C.D., Faulds, J.E., and Hinz, N.H., 2005, The upper reaches of the Sierra Nevada auriferous gold channels: Geological Society of Nevada Symposium (in press).

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Reviews of Earth and Planetary Sciences, v. 32, p. 67–89, doi: 10.1146/ annurev.earth.32.091003.143456. Oldow, J.S., 1984, Evolution of a late Mesozoic back-arc fold and thrust belt, northwestern Great Basin, U.S.A.: Tectonophysics, v. 102, p. 245–274, doi: 10.1016/0040-1951(84)90016-7. Oldow, J.S., 1992, Late Cenozoic displacement partitioning in the northwestern Great Basin, in Stewart, J., ed., Structure, tectonics and mineralization of the Walker Lane—Walker Lane Symposium Proceedings: Reno, Geological Society of Nevada, p. 17–52. Oldow, J.S., Aiken, C.L.V., Hare, J.L., Ferguson, J.F., and Hardyman, R.F., 2001, Active displacement transfer and differential block motion within the central Walker Lane, western Great Basin: Geology, v. 29, p. 19–22, doi: 10.1130/0091-7613(2001)029<0019:ADTADB>2.0.CO;2. Park, H., Trexler, J., Cashman, P., and Mass, K.B., 2004, Local initiation of Walker Lane tectonism prior to 3.6 Ma recorded in Neogene sediments at Honey Lake basin, northeastern California: Geological Society of America Abstracts with Programs, v. 36, no. 4, p. 16–17. Petit, J.P., 1987, Criteria for the sense of movement on fault surfaces in brittle rocks: Journal of Structural Geology, v. 9, p. 597–608, doi: 10.1016/01918141(87)90145-3. Sanders, C.O., and Slemmons, D.B., 1979, Recent crustal movements in the central Sierra Nevada—Walker Lane region of California-Nevada: Part III, the Olinghouse fault zone: Tectonophysics, v. 52, p. 585–597, doi: 10.1016/0040-1951(79)90273-7. Saucedo, G.J., and Wagner, S.L., 1992, Geologic map of the Chico quadrangle: California Division of Mines and Geology, Regional Geologic Map Series Map No. 7A, scale 1:250,000. Stewart, J.H., 1988, Tectonics of the Walker Lane belt, western Great Basin: Mesozoic and Cenozoic deformation in a zone of shear, in Ernst, W.G., ed., The geotectonic development of California: Englewood Cliffs, New Jersey, Prentice Hall, p. 7186. Surpless, B.E., Stockli, D.F., Dumitru, T.A., and Miller, E.L., 2002, Two-phase westward encroachment of Basin and Range extension into the northern Sierra Nevada: Tectonics, v. 21, p. 2-1–2-13.

Taymaz, T.J., Jackson, J., and McKenzie, D., 1991, Active tectonics of the north and central Aegean Sea: International Geophysical Journal, v. 106, p. 433–490. 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. Trexler, J.H., Cashman, P.H., Henry, C.D., Muntean, T., Schwartz, K., TenBrink, A., Faulds, J.E., Perkins, M., and Kelly, T., 2000, Neogene basins in western Nevada document the tectonic history of the Sierra Nevada—Basin and Range transition zone for the last 12 Ma, in Lageson, D.R., Peters, S.G., and Lahren, M.M., eds., Great Basin and Sierra Nevada: Boulder, Colorado, Geological Society of America Field Guide 2, p. 97–116. U.S. Geological Survey, 2004, Quaternary fault and fold data-base of the United States: U.S. Geological Survey, http://qfaults.cr.usgs.gov/index.htm. Wagner, D.L., Renne, P.R., Deino, A.L., and Saucedo, G.J., 2000, Age and origin of the Lovejoy Basalt of northern California: Geological Society of America Abstracts with Programs, v. 32, no. 7, p. A159. Wallace, A.B., 1975, Geology and mineral deposits of the Pyramid District, Washoe County, Nevada [Ph.D. dissertation]: Reno, University of Nevada, 162 p. Wernicke, B., 1992, Cenozoic extensional tectonics of the U.S. Cordillera, in Burchfiel, B.C., Lipman, P.W. and Zoback, M.L., eds., The Cordilleran orogen: Conterminous U.S.: Boulder, Geological Society of America, The Geology of North America, v. G-3, p. 553–581. Wernicke, B.P., Clayton, R.W., Ducea, M.N., Jones, C.H., Park, S.K., Ruppert, S.D., Saleeby, J.B., Snow, J.K., Squires, L.J., Fliedner, M.M., Jiracek, G.R., Keller, G.R., Klemperer, S.L., Luetgert, J.H., Malin, P.E., Miller, K.C., Mooney, W.D., Oliver, H.W., Phinney, R.A., 1996, Origin of high mountains in the continents: The southern Sierra Nevada: Science, v. 271, p. 190–193. Wilcox, R.E., Harding, T.P., and Seely, D.R., 1973, Basic wrench tectonics: AAPG Bulletin, v. 57, p. 74–96. Wills, C.J., and Borchardt, G., 1993, Holocene slip rate and earthquake recurrence on the Honey Lake fault zone, northeastern California: Geology, v. 21, p. 853–856, doi: 10.1130/0091-7613(1993)021<0853:HSRAER>2.3.CO;2.

Printed in the USA

Geological Society of America Field Guide 6 2005

Brittle deformation, fluid flow, and diagenesis in sandstone at Valley of Fire State Park, Nevada Peter Eichhubl Texas A&M University–Corpus Christi, Corpus Christi, Texas 78412, USA Eric Flodin Indiana University–Purdue University, Fort Wayne, Indiana 46805, USA

ABSTRACT The interaction among brittle deformation, fluid flow, and diagenesis is displayed at Valley of Fire, southern Nevada, where diagenetic iron oxide and hydroxide stains provided a visible record of paleofluid flow in Jurassic Aztec Sandstone. Deformation features include deformation bands, joints, and faults composed of deformation bands and sheared joints. Faults formed by shear along joints, formation of splay fractures, and linkage of fault segments. Measurements of fault permeability, combined with numerical permeability upscaling, indicate that these faults impede cross-fault fluid flow, with cross-fault permeability reduced by two orders of magnitude relative to the host sandstone, whereas fault-parallel permeability is enhanced by nearly one order of magnitude. A reconstruction of paleofluid flow in the Aztec Sandstone is based on detailed mapping of multicolored alteration patterns and their cross-cutting relations with brittle structures. These patterns resulted from syndepositional reddening of the eolian sandstone and repeated episodes of dissolution, mobilization, and reprecipitation of iron oxide and hydroxide. The distribution of alteration patterns indicates that regional-scale fluid migration pathways were controlled by stratigraphic contacts, thrust faults, and high-angle oblique-slip faults. Outcrop-scale focusing of fluid flow was controlled by structural heterogeneities such as joints, joint-based faults, and deformation bands as well as the sedimentary architecture. The complex interaction of structural heterogeneities with alteration in this exhumed analog of a fractured and faulted sandstone aquifer is consistent with their measured hydraulic properties demonstrating the significance of structural heterogeneities for focused fluid flow in porous sandstone aquifers. Keywords: fluid flow, sandstone, diagenesis, faults and faulting, jointing, permeability. INTRODUCTION Brittle deformation, fluid flow, and diagenetic processes are coupled phenomena (Dickinson and Milliken, 1995; Eichhubl and Behl, 1998; Laubach et al., 2004; Eichhubl et al., 2004; Gale et

al., 2004). Brittle structures such as joints, deformation bands, and faults composed of joints and deformation bands affect fluid flow and thus mass transfer and chemical alteration in the subsurface (Antonellini and Aydin, 1994; Eichhubl and Boles, 2000; Flodin et al., 2005). Diagenetic processes such as compaction, grain dis-

Eichhubl, P., and Flodin, E., 2005, Brittle deformation, fluid flow, and diagenesis in sandstone at Valley of Fire State Park, Nevada, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 151–167, doi: 10.1130/2005.fld006(07). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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solution, and cementation affect the physical properties of rock and thus the deformation behavior (Gross, 1995; Flodin et al., 2003). This field trip examines the interaction among these processes in an exhumed sandstone aquifer at Valley of Fire State Park, located in the Northern Muddy Mountains along the shores of Lake Mead, ~60 km northwest of Las Vegas, Nevada (Fig. 1). The Northern Muddy Mountains are part of the Basin and Range physiographic province, located ~50 km west of the western edge of the Colorado Plateau. The park also contains a frontal section of the Sevier orogenic belt of Cretaceous age. This trip is dedicated to the interplay among depositional architecture, Cretaceous and Tertiary deformation, fluid flow, and chemical mass transfer. STRATIGRAPHY OF VALLEY OF FIRE The name of the park relates to the red, purple, orange, yellow, and white hues of the Upper Triassic and Jurassic Aztec Sandstone (Bohannon, 1977, 1983). This eolian sandstone was deposited in a backarc setting (Marzolf, 1983, 1990) and is considered equivalent to the Navajo Sandstone of the Colorado Plateau and the Nugget Sandstone in the Wyoming thrust belt (Poole, 1964; Blakey, 1989). The Aztec Sandstone at Valley of Fire is part of a Mesozoic clastic sequence that overlies Upper Paleozoic carbonates and shale (Bohannon, 1983). Triassic redbeds of the Moenkopi, Chinle, and Moenave Formations, with a combined thickness of up to 2100 m (Bohannon, 1977), are composed of sand-, silt-, and mudstone with evaporite layers (Marzolf, 1990). The Aztec Sandstone is unconformably overlain, with a discordance of locally up to 10°, by a distinctive conglomerate that forms the basal member of the Cretaceous Willow

Tank Formation. The Willow Tank Formation, composed of predominantly mudstone, and the overlying Baseline Sandstone have a combined thickness of ~1300 m (Bohannon, 1977). These units are interpreted as synorogenic foreland deposits of the eastward-directed Cretaceous Sevier thrust sheets (Bohannon, 1983). The basal conglomerate of the Willow Tank Formation includes locally derived red Aztec sandstone components (Longwell, 1949), indicative of partial exhumation and erosion of the Aztec Sandstone subsequent to its deposition (Bohannon, 1983). The Willow Tank Formation and lower portions of the Cretaceous Baseline Sandstone are overthrust by Aztec Sandstone of the Willow Tank thrust sheet (Fig. 2), the lowest of the Sevier thrust sheets in Valley of Fire (Longwell, 1949; Bohannon, 1983). Tearfault Mesa in the northern part of Figure 2 forms a klippe of the Willow Tank thrust sheet. Cretaceous strata are steepened and locally overturned along the eastern margin of this klippe (Longwell, 1949), indicating that the eastern edge of the klippe represents the leading edge of the thrust sheet. Upper portions of the Baseline Sandstone are deposited on top of the Willow Tank thrust sheet, constraining the age of thrusting to Late Cretaceous (Maastrichtian) (Longwell, 1949; Bohannon, 1983). After a hiatus during Paleogene times, conformable deposition of clastic units resumed during Oligocene times (Bohannon, 1983). A ~25° angular unconformity at the base of the upper Miocene (10–4 Ma) Muddy Creek Formation (Bohannon et al., 1993) constrains the time of regional tilting of the Aztec Sandstone, presently dipping 20–30° NE in the area visited on this trip (Fig. 2). The Muddy Creek Formation, in turn, dips 5–8° NE, indicating continued regional tilting throughout the Late Miocene. The Aztec Sandstone and overlying Cretaceous strata are folded into a gentle NE-plunging syncline referred to as Overton syncline by Carpenter and Carpenter (1994) (Fig. 2). DEFORMATION OF THE AZTEC SANDSTONE

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Figure 1. Location of map of Valley of Fire State Park in southern Nevada.

Deformation structures in the Aztec Sandstone include deformation bands, joints and sheared joints, and faults that are composed of deformation bands, slip surfaces, sheared joints, and breccia zones. Deformation bands are tabular zones of localized deformation that are frequently more resistant to erosion. Thus, they are often observed on the outcrop as ridges ranging in thickness from 1 mm to 10 cm, with a typical width of ~1 cm (Antonellini and Aydin, 1994). At Valley of Fire, Hill (1989) distinguished three earlier sets of deformation bands, striking northwest, north-northwest, and north-northeast, that are characterized by predominant band-parallel compaction. Following Mollema and Antonellini (1996) and Sternlof and Pollard (2001), we refer to these deformation bands as compaction bands. Repeated mutual cross-cutting suggests that these three sets formed concurrently. These three sets of compaction bands are cross-cut by three sets of deformation bands (Flodin and Aydin, 2004) that exhibit macroscopic shear offsets and are therefore referred to as shear bands. One set of shear bands, with slip ranging from 1 to 3 cm, parallels

Brittle deformation, fluid flow, and diagenesis in sandstone 114°32'

4042000 36°30'

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Tertiary & Quaternary Cretaceous Jurassic Aztec Sandstone upper red and banded red and white units middle alteration units lower red unit

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Figure 2. Geologic map of the northern part of Valley of Fire State Park. WP—Waterpocket fault; BH—Bighorn fault; W—Wall fault; L—Lonewolf fault; C—Classic fault; M—Mouses Tank fault; B—Baseline fault. Figure from Eichhubl et al. (2004).

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depositional boundaries of the cross-bedded strata (Hill, 1989). The other two sets of shear bands usually occur as subvertical zones of multiple shear bands and associated slip surfaces, with slip on the order of 1–10 cm (Flodin and Aydin, 2004). The topto-the-east sense of shear of the bed-parallel shear bands is kinematically consistent with Sevier thrusting, suggesting that these shear bands formed concurrently with thrusting (Hill, 1989). Taylor et al. (1999) distinguished four sets of joints in the Aztec Sandstone. The first set is composed of parallel vertical joints that strike roughly north-south. This set is locally replaced by two sets of joints that also strike north-south but form two intersecting sets with an acute angle of ~30°. Joints of the fourth set are slightly sinuous and do not show a strong preferred orientation, but may favor a roughly east-west orientation. Joints have been observed to consistently crosscut compaction and shear bands and are thus interpreted to form after deformation bands and concurrently with, and possibly also prior to, strike-slip faulting (Myers, 1999; Flodin and Aydin, 2004). Two sets of oblique-slip faults with predominant strike-slip and lesser normal-slip components occur as two sets: a left-lateral fault set striking NNE, and a right-lateral set striking NW (Myers, 1999; Taylor, 1999; Flodin and Aydin, 2004) (Fig. 2). The NNE-striking set is more prominently developed on the regional scale, but both sets show mutually abutting geometries on a local scale, suggesting that both sets were active concurrently (Flodin and Aydin, 2004). This network of faults within the Valley of Fire is bounded by the Waterpocket fault system to the west and the Baseline fault system to the east (Fig. 2). Both the Waterpocket and Baseline fault systems show nearly 2.5 km of apparent left-lateral separation. They appear to be part of a larger family of ~N-trending left-lateral faults that occur to the northeast along Weiser Ridge and to the north in the southern and eastern Mormon Mountains and the Tule Spring Hills (Bohannon, 1992; Anderson and Barnhard, 1993; Axen, 1993). These faults offset the Aztec–Willow Tank formation contact. Following the maps of Bohannon (1977) and Carpenter (1989), the youngest strata offset by faults of the Baseline fault system are the Miocene Horse Spring Formation and probably lower sections of the Upper Miocene Muddy Creek Formation (10–4 Ma; Bohannon et al., 1993). These faults are thus considered to be associated with Basin and Range tectonics (Flodin and Aydin, 2004). DIAGENESIS OF AZTEC SANDSTONE The Aztec Sandstone is a fine- to medium-grained subarkose (Marzolf, 1983) with up to 8% feldspar (Flodin et al., 2003) and a smaller amount of lithic components. The sandstone is generally friable with a porosity of 15%–25% and permeability of 100– 2500 md (Flodin et al., 2003). Grains are subrounded to rounded, characteristic of the eolian depositional environment. Grains are weakly indented throughout the section as a result of pressure solution, and pore space is partially filled with clay cement, predominantly kaolinite and lesser mixed-layer illite-smectite. The mixed-layer illite-smectite is 90% illite with a Reichweite of 3.

In the stratigraphically lowest part of the Aztec Sandstone, illite replaces the mixed-layer illite-smectite (Eichhubl et al., 2004). Minor amounts of quartz overgrowth cement were observed in stratigraphically lower parts of the Aztec Sandstone. Quartz cement was also found locally within the damage zone and to both sides of the Bighorn fault (Fig. 2), cementing allochthonous Aztec Sandstone of the Willow Tank thrust sheet as well as fault rock that is part of the autochthonous Aztec Sandstone. We estimate that the maximum burial depth at the top of the Aztec Sandstone corresponds closely to the thickness of the post-Jurassic units to the east, amounting to ~1600 m (based on stratigraphic thicknesses provided by Bohannon, 1983). The occurrence of quartz overgrowth cement and diagenetic illite in the lowermost part of the Aztec Sandstone, typically requiring burial temperatures in excess of ~80 °C (Worden and Morad, 2000, 2003), appears in agreement with this estimate. The geothermal gradient could have been elevated along the thrust front due to fluid flow, consistent with the occurrence of quartz cement along and below the Willow Tank thrust sheet. The characteristic hues of red, orange, purple, yellow, and white of the Aztec Sandstone result from varying amounts and forms of iron oxide and hydroxide cement, the dominant pigments being hematite and goethite. The uniform red color of the lower red and upper red alteration units of the Aztec Sandstone results from thin grain coats of hematite. In thin section, hematite forms a mottled brownish coloration of grain surfaces. In samples of lower red sandstone containing quartz overgrowth cement, the hematite coat predates the quartz overgrowth, indicative of an early diagenetic or syndepositional origin of the hematite coat. Compared to grain coats in red sandstone, grain coats in yellow sandstone (Munsell 10YR 7/6 yellow to 10YR 8/2 very pale brown; Munsell Soil Color Charts, 1994) are more patchy and partly recrystallized to ≤1-μm-sized crystals. Based on X-ray diffractograms, the dominant Fe-oxide is goethite. In orange-colored sandstone (Munsell 2.5 YR 5/8 red), hematite forms 5–10-μm-sized globules that are attached to grain surfaces. The equally spaced distribution of hematite globules along grain surfaces is suggestive of globule formation by local dissolution of earlier grain coats and reprecipitation of hematite with accompanying crystal coarsening. Purple coloration (Rock Color Chart 5RP 6/2 pale red purple to 5RP 4/2 grayish red purple, Rock-Color Chart Committee, 1970; and Munsell 10R 6/3 pale red) is caused by 1–3-μm-sized grains of Fe-oxide, identified by X-ray diffraction to be dominantly goethite. Some samples of purple and yellow sandstone also contain smaller amounts of the sulfates alunite KAl3(SO4)2(OH)6 and jarosite KFe3(SO4)2(OH)6. Jarosite provides a mottled brownish color to the otherwise uniform yellow sandstone. White sandstone (Munsell 10YR 8/1–7.5YR 8/1 white) is devoid of grain coats but contains sparse flakes of coarse-crystalline hematite. IMPACT OF STRUCTURES ON FLUID FLOW The flow characteristics of faults formed by shearing of joints in Aztec Sandstone, Valley of Fire, have been discussed by Myers

Brittle deformation, fluid flow, and diagenesis in sandstone (1999), Taylor (1999), Aydin (2000), Flodin et al. (2001), Jourde et al. (2002), Flodin and Aydin (2004), and Flodin et al. (2005). The faults impart a dual influence on flow behavior, where finegrained fault rock or gouge at the core of the fault zone reduces cross-fault permeability, and joints and slightly sheared joints in the surrounding fault-damage zone increase along-fault permeability. These flow properties correspond to a combined barrierconduit system according to Caine et al. (1996). Considering cross-fault permeability, Flodin et al. (2005) described petrophysical properties for a suite of fault rocks. They found an average permeability reduction of three orders of magnitude in fault-rock samples with respect to host-rock samples and that overall cross-fault permeability decreases with increasing average shear strain. Taylor et al. (1999) used numerical simulations to calculate flow fields perturbed by the presence of joints and sheared joints. By comparing their numerical results with field relationships, they estimated permeabilities for joints and slightly sheared joints in the Valley of Fire to be about five orders of magnitude greater than hostrock permeability. Myers (1999), using power-averaging techniques, calculated upscaled permeabilities for a series of outcrop-scale fault maps. He found significant permeability anisotropy with ratios between alongand cross-fault permeability ranging from one to four orders of magnitude. Jourde et al. (2002) took a numerical approach to the same upscaling problem and demonstrated a reduction of faultnormal permeability by two orders of magnitude relative to the host rock permeability. Permeability parallel to the fault, on the other hand, was found to be increased by nearly one order of magnitude relative to the host rock due to the presence of connected joint networks with a preferred orientation parallel to the fault. The flow characteristics of individual compaction bands in the Valley of Fire were addressed by Taylor and Pollard (2000). At a larger-scale of observation, Sternlof et al. (2004) investigated the combined flow effects of arrays of compaction bands. They considered three different characteristic deformation band array geometries—parallel, cross-hatch, and anastomosing—and concluded that all were capable of imparting up to two orders of magnitude permeability reduction and a corresponding increase in permeability anisotropy. INTERACTIONS AMONG STRUCTURAL, FLUID FLOW, AND DIAGENETIC PROCESSES Rock units of common alteration color were mapped (Eichhubl et al., 2004) as shown in Figure 3, distinguishing nine types of alteration based either on the dominant color (red, yellow, etc.) or a combination of two colors (e.g., banded orange-and-white, banded purple-and-yellow). Rock units of common alteration color will be addressed in the field as alteration bands or alteration units. A schematic composite cross section across the diagenetic alteration units of the Aztec Sandstone (Fig. 4) illustrates the upward succession of a lower red unit, overlain by middle purple, yellow, orange, and white alteration units, and capped by an upper red alteration unit. Although these alteration units

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roughly parallel the stratigraphy, we will see at several stops that alteration bands and contacts cut depositional boundaries and thus postdate deposition. These alteration bands and units are thus the result of diagenetic processes. Based on diagenetic and structural cross-cutting relations, Eichhubl et al. (2004) distinguished three stages of alteration (Fig. 5): a first stage of syndepositional reddening of the sand or sandstone and two stages of dissolution or bleaching, iron remobilization, and reprecipitation of iron oxide and hydroxide. The first stage of bleaching and iron remobilization, precipitating dominantly hematite, is attributed to upward migration of reducing basinal fluid during and subsequent to Late Cretaceous Sevier thrusting and foreland deposition of clastic sediments. A second stage of bleaching and iron remobilization, precipitating predominantly goethite and minor iron sulfates, occurred during Miocene strike-slip faulting associated with Basin and Range tectonics. This second stage is explained by mixing of reducing sulfide-rich basinal fluid with meteoric water entering the aquifer. By relating patterns of alteration to fluid flow, we can derive that regional-scale fluid migration pathways were largely controlled by stratigraphic contacts, thrust faults, and high-angle oblique-slip faults. The outcrop-scale focusing of flow was controlled by structural heterogeneities such as joints, joint-based faults, and deformation bands as well as the sedimentary architecture. Deformation bands tend to form alteration boundaries, suggesting they acted as baffles to subsurface fluid flow. Joints and sheared joints are frequently surrounded by halos of enhanced reddening or bleaching indicative of focused fluid flow along joints. The interaction of structural heterogeneities with paleofluid flow as inferred using the diagenetic alteration record is consistent with their measured hydraulic properties and flow models. Taylor et al. (1999) compared the distribution of alteration halos around isolated and interconnected joints with computed flow patterns, emphasizing the importance of joint connectivity for fluid flow in sandstone of high matrix permeability such as the Aztec Sandstone. The combined field, laboratory, and modeling evidence demonstrates the significance of structural heterogeneities for controlling fluid flow properties of porous sandstone aquifers. In addition to the depositional heterogeneity, structural features focus fluid flow and chemical alteration. On a regional scale, the diagenetic and structural record at Valley of Fire illustrates the coupling of tectonic, fluid flow, and alteration processes. DESCRIPTION OF STOPS All stops described here are located within Valley of Fire State Park, within a driving distance of 5.5 mi (9 km) from each other (Fig. 6). Stops 2–6 require some hiking over easy terrain, with hikes not exceeding 1.5 mi (2.4 km) roundtrip. Although described here as a continuous trip, a visit of all stops will require one and a half to two full days. Best times to visit are during spring and fall; average daytime highs exceed 105 °F (40 °C) in the months of July and August.

P. Eichhubl and E. Flodin

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Figure 3. Map of diagenetic alteration units for area shown in Figure 2.

Brittle deformation, fluid flow, and diagenesis in sandstone Southwest Willow Tank Thrust

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Directions to Stop 1 Valley of Fire State Park is accessed most conveniently from Las Vegas. Leave Las Vegas heading north on I-15, take Exit 75 at Crystal, and proceed on Nevada Route 169 southeast to Valley of Fire State Park. In the park, 18 mi (29 km) from I-15 Exit 75, turn left and park at the visitor’s center (Stop 1). Park entrance fees are collected. Please note that sample collection is not permitted in Valley of Fire State Park without prior permission by the park administration. Parking is limited to designated parking areas.

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Figure 4. Schematic composite cross section across diagenetic alteration units in Figure 3, based on field sketches taken from several vantage points. Due to the northwestward plunge of the alteration units, the top parts of the section reflect observations in the northern portion of the map area, and lower parts are taken from observations in the southern portion of the map. Because alteration features do not project along strike over the extent of the map area, this section represents an idealized composite. Although unit thicknesses and inclination are approximately to scale, horizontal distances are condensed to about half of the vertical scale.

1 km

The road enters a valley that follows the N-S–striking Mouses Tank fault. Continue past the Mouses Tank parking area and park at Rainbow Vista, 1.8 mi (2.9 km) from the visitor’s center. Stop 2a—Rainbow Vista Following the road leading up from the visitor’s center, we traversed the stratigraphically lower 800–900-m-thick section of the Aztec Sandstone. Apart from joint surfaces covered by darkbrown varnish, this section is characterized by a red stain that

Stop 1—Valley of Fire State Park Visitor’s Center The visitor’s center provides a panoramic overlook of Valley of Fire Wash and the Muddy Mountains to the south. The wash follows the outcrop of less resistant Triassic sand- and siltstone units that underlie the cliff-forming Aztec Sandstone behind, and to the north of, the visitor’s center. The visitor’s center is situated on reddish sandstone and siltstone of the Upper Triassic Moenave Formation, with older Triassic Chinle and Moenkopi Formations underlying the central part of the wash. Red Aztec Sandstone is exposed in the lower parts of the Muddy Mountains to the south, including the Fire Alcove just to the south of our present position, as well as in the low hills in the upper part of Valley of Fire Wash (Atlatl Rock). The upper section of the Muddy Mountains is composed of Paleozoic limestone and dolostone, separated from the Mesozoic clastic units by the E-W–striking Arrowhead fault (Longwell, 1949). Directions to Stop 2 Continue along the road that leads from the visitor’s center into the northern part of the park. As the road climbs, we traverse from the Triassic red beds into the basal section of the Aztec Sandstone.

A. Late Jurassic

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meteoric fluid

basinal fluid

Figure 5. Stages of diagenetic alteration, fluid flow, and deformation in the Aztec Sandstone at Valley of Fire. (A) Syndepositional reddening of eolian sandstone. (B) First stage of bleaching and iron remobilization in upper section of the Aztec Sandstone associated with Sevier thrusting. Field observations suggest local flow along strike of the orogenic front, but a regional direction of flow away from the orogen is inferred. (C) Second stage of bleaching and iron remobilization associated with mixing of basinal and meteoric fluids during Basin and Range tectonics.

P. Eichhubl and E. Flodin

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Figure 6. Map of field trip stops in Valley of Fire State Park; same area as Figure 2.

is uniform without regard to grain size variations among crossstratified layers. Immediately to the north of the Rainbow Vista parking area, the color changes to yellow and purple over a relatively sharp 1–5-cm-wide boundary. This color boundary is offset along the Mouses Tank fault by 150 m of left-lateral slip. The boundary is parallel to bedding over distances of 10–100 m but

frequently climbs, and occasionally drops, across bedding over distances of 10–50 m. This color boundary is thus diagenetic and not depositional in origin. This and other similar color boundaries will thus be referred to as alteration contact. In the Rainbow Vista area, the average orientation of the alteration contact dips to the northeast by up to 49° compared to bedding at 22°. The uni-

Brittle deformation, fluid flow, and diagenesis in sandstone formly red sandstone in the lower part of the Aztec Sandstone is referred to as “lower red unit” in Figures 2 and 3, and the yellow and purple alteration forms part of the middle alteration units. The alteration contact has a lobate-cuspate geometry, shown schematically in Figure 4, with lobes of purple or yellow sandstone penetrating the lower red sandstone. One such lobe is visible on the cliff facing Rainbow Vista, west of the road. We interpret this lobate-cuspate alteration contact as a reaction boundary that migrated into the red sandstone, changing the red to yellow. This interpretation is consistent with a syndepositional origin of the hematite grain coat in uniformly red sandstone. The structural offset of this alteration contact along the Mouses Tank fault indicates that this bleaching occurred prior to slip on the Mouses Tank fault, or at least prior to the last 150 m of slip along this fault. We will see evidence at subsequent stops that the syndepositional reddening of the Aztec Sandstone and the yellow and purple alteration are separated by an intermediate alteration stage characterized by banded orange-and-white rock colors. This second stage of alteration has been completely overprinted at Rainbow Vista by the third stage of yellow and purple alteration. A flat exposure of Aztec Sandstone ~20 m east of the parking area exhibits cross-cutting relations among compaction bands, shear bands, joints, and sheared joints. Subvertical, NNW-SSE–striking, 1–2-cm-thick compaction bands are the oldest structures. They are offset by ~1 cm of slip along subhorizontal shear bands and accompanying slip surfaces that follow depositional surfaces. Slip along these shear bands is top to the south. Joints follow earlier compaction and shear bands. The youngest discernible stage of deformation is slip along these joints, resulting in the formation of tail cracks at a high angle to the sheared joints. Where compaction bands cross the alteration contact between lower red and middle yellow and purple sandstone, the alteration contact appears in some instances offset along the bands even where depositional surfaces are not offset. These offsets of the alteration contacts are interpreted to reflect the retardation of the moving reaction boundary across the lowerpermeability compaction band. Examples are visible on the cliffs to the east of Rainbow Vista. We will discuss this effect again at Stop 6a. Cross the park road and walk 60 m W to exposures of the N-S–striking Mouses Tank fault. Stop 2b—Mouses Tank Fault as a Joint-Based Fault This stop provides the first example of a fault zone that formed by shearing along preexisting joint zones. This zone has at least 150 m of net left-lateral strike-slip and is structurally quite complex (Fig. 7A). The 0.5–1-m-thick fault core is composed of fault rock and slip surfaces, the latter with an overall orientation parallel to the fault strike. The 2–3-m-thick fault damage zone is composed of joints, sheared joints, and deformation bands. Joints are frequently oriented at an oblique angle to sheared joints, with a characteristic splay fracture geometry.

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Following Myers (1999) and Flodin and Aydin (2004), the evolution of these faults includes the formation of joints, slip along joints, formation of splay or tail fractures, and linkage of sheared joints. Structural elements that comprise joint-based faults include joints, sheared joints, shear bands, fragmented rock, fault rock, and slip surfaces. Four end members of preexisting joint zone geometries are recognized (Myers and Aydin, 2004; Flodin and Aydin, 2004): (1) en-échelon joint zones that have step-sense opposite to shear-sense (e.g., right-stepping and left-lateral shearing); (2) en-échelon joint zones that have step-sense similar to shear-sense (e.g., right-stepping and right-lateral shearing); (3) subparallel joint zones characterized by a large joint-length to joint-spacing ratio; and (4) individual joints with attendant fringe joints (or hackle). For en-échelon joint zones that have a step-sense opposite to the slip-sense, the overlapping region between stepping joints is subject to a localized contractional strain (Segall and Pollard, 1980; Lin and Logan, 1991), which results in the frictional breakdown of host rock material. Damage in the form of joints and sheared joints is outwardly developed (Myers, 1999). In contrast, enéchelon joint zones that have the same step-sense and slipsense are subject to a localized dilational strain that results in the fragmentation of rock that spans the overlapping en-échelon joints by inwardly directed splay fractures. For subparallel joint zones, strain is accommodated by the formation of splay fractures that span the distance between the overlapping joint segments. Accumulated shear strain is preferentially localized along the closely spaced subparallel sheared joints (Myers and Aydin, 2004). This stop also illustrates large-scale permeability effects of the fault zones. From a hydrologic perspective, the joint-based faults in the Valley of Fire are composed of high permeability components (joints, splay fractures, and slip surfaces) and low permeability components (fault rock, deformation bands, and sheared joints) embedded in a matrix with intermediate permeability. Using a detailed map of a portion of Mouses Tank fault drafted by Myers (1999), Jourde et al. (2002) derived bulk permeability values for this fault zone using numerical upscaling methods (Fig. 7). The upscaling work flow starts with subcentimeter-resolution field maps that distinguish the various elements of the fault zone (Fig. 7A). The permeability values of each fine scale fault zone element (joints, sheared joints, fault related deformation bands, slip surfaces, fault rock, and the matrix rock) are either measured or estimated (Fig. 7B) and then input into the detailed description (Fig. 7C). Numerical simulation of the finescale input map yields the larger-scale permeability of the fault zone of interest (Fig. 7D). For both map regions featured in Figure 7, the fault-perpendicular component of permeability (k1) is reduced by over two orders of magnitude relative to the host rock. This large reduction is clearly due to the wide zone of fault rock and to the fact that there are no slip surfaces traversing this zone in the perpendicular direction. The dense regions of deformation bands at stepovers and sheared joints emanating out from the gouge also contribute

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Figure 7. Workflow used to upscale fault zone permeability. (A) Map of Mouses Tank fault with 150 m of strike-slip, from Myers (1999). Values of large-scale permeability k1 and k2 (in md) as calculated by Jourde et al. (2002) for each map area as computed in step (D). (B) Laboratory analysis for fault element permeability. (C) Rasterization of field map and assignment of element permeability values. (D) Numerical calculation of large-scale permeability.

k2

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to the low fault-normal permeability. The continuous slip surfaces in the direction along the fault lead to enhanced permeability in the fault-parallel direction. Directions to Stop 3 Follow the park road north toward White Dome for 1 mile (1.6 km) and park at the next parking area (parking area 1) to the left of the road. Walk ~120 m west to a prominent south-facing fault exposure. Stop 3a—Structure, Petrophysics, and Permeability of a Joint-Based Fault with ~25 m Slip This location provides a well-exposed cross section across the Lonewolf fault with ~25 m of left-lateral strike-slip (Fig. 2) and associated secondary and higher order structures. This exposure was one of three sampling stations in a study of

slip surface

fault rock petrophysics (Flodin et al., 2005). Four basic rock elements are identified (Fig. 8): (1) undamaged host rock; (2) damaged host rock; (3) fragmented rock; and (4) fault rock. No bulk mineralogical changes due to fault zone cementation or mineral alteration are detected when comparing host and fault rock. Fault rock permeability is one to three orders of magnitude lower than median host rock permeability. Porosity reductions are less pronounced and show considerable overlap in values between the sample suites. Some fragmented rock samples appear to have dilated with respect to median host rock porosity. Median grain sizes for fault rock samples range from 3 to 51 μm, which is up to two orders of magnitude reduction from host rock median grain sizes. There appears to be a lower limit of median grain size of 3 μm for fault rock samples irrespective of average fault shear strain. Fault rock capillary injection pressures range from one to almost two orders of magnitude higher than host rock equivalent. Based on these measurements, we conclude that faults formed by shearing of

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Figure 8. Backscatter electron images of deformed Aztec Sandstone at Stop 3a. White minerals are potassium feldspar; light gray minerals are quartz; dark gray minerals are clays; black is pore space. (A) Host rock fragment showing earliest signs of grain-scale fragmentation. Nearly every grain shows varying degrees of fracturing; however, the primary pore space is preserved. f—potassium feldspar, p—pore space, q—quartz. (B) Host rock collected adjacent to fault core. Note that many of the grains are fractured and that much of the pore space is occupied by clay minerals (c). (C) Host rock collected adjacent to the fault core showing signs of severe grain-scale fragmentation, collapse of pores, and low primary porosity. (D) Well-developed fault rock characterized by severe grain size reduction and a complete loss of primary porosity. Dashed lines are approximate locations of slip surfaces. Note the variability with respect to pore space (black) between the left and right sides of the image. Figure adapted from Flodin et al. (2005).

joints in high permeability sandstone systems will act as significant barriers to fluid flow on short time scales and might be capable of sealing small to moderate hydrocarbon columns on a geologic time scale as well, assuming adequate continuity of the fault rock over large areas of the fault. Follow the fault NNE across the park road. Walk E along the far side of the park road for ~20 m, then veer to the NW at the first small valley and walk ~5 m. Stop 3b—Small-Scale Material Rotation Material rotations play an important role in the conceptual framework for the development of the fault system in the Valley of Fire. Importantly, material rotations allow for newly resolved shear stresses across fractures that originated as opening-mode joints. At this stop we see evidence for small-scale material rotation in the vicinity of a fault zone. This example is from a left stepover between right-lateral bounding faults that show

cumulative separations on the order of a few meters. Traversing the stepover area are a set of left-lateral faults and two steeply dipping, preexisting deformation bands that enter the fault zone from the northeast. Outside of the fault zone, both deformation bands continue with a straight trace for many meters toward the northwest. Within the fault zone, the deformation bands are broken and offset by left-lateral faults that are bound between the two overlapping right-lateral faults. Between the left-lateral faults, the deformation band segments approximately retain their straight trace. The deformation band segments are progressively rotated in a clockwise sense with increasing distance to the lower right-lateral fault. At the last documented juncture between the deformation bands and the right-lateral fault, the bands are rotated by as much as 60°. A best-fit great circle through the orientation data indicates the deformation bands were rotated about axes that plunge ~70°. Follow the main trace of the Lonewolf fault zone 220 m to the NNE.

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Stop 3c—Fault Hierarchies and Network Evolution Mutual abutting relationships between different sets of leftand right-lateral faults and splay fractures at this locality imply a hierarchical sequence of formation. Here, the largest structure is the left-lateral Lonewolf fault. We consider the Lonewolf fault primary in nature because it bounds all other structures. Emanating from the Lonewolf fault are splay fractures and rightlateral faults that share the same orientation. These structures are considered to be second generation because they appear to have formed in response to activity along the primary Lonewolf fault. All secondary structures have at least one end that abuts against a primary left-lateral fault. Two of the second-generation rightlateral faults extend through the map area to the southwest where they join with the end of the Classic fault to the southwest. Emanating from, and localized between, these two right-lateral faults are splay fractures and left-lateral faults, both of which have the same intersection angle. These structures are considered to be third-generation because they appear to be genetically related to activity along the second-generation structures. All thirdgeneration structures are bound on at least one end by a second generation right-lateral fault. Further branching occurs from the third-generation left-lateral faults in the form of fourth-generation splay fractures, the subsequent shearing of which produced a fourth-generation of right-lateral faults. The highest order structure identified in this area is a fifth-generation left-lateral fault and related splay fractures. In the conceptual model shown in Figure 9, the first generation of faults with a left-lateral sense of shearing produces opening-mode splay fractures that are primarily localized at or near the

B

A

C

end, and in the immediate vicinity, of the first-generation structures. However, due to mechanical interaction between adjacent first-generation faults (Martel, 1990) and in response to increasing slip, some splay fractures propagate across the distance of undeformed rock that spans these neighboring structures (Fig. 9). Due to local stress perturbations between overlapping faults (e.g., Segall and Pollard, 1980; Kattenhorn et al., 2000; Bourne and Willemse, 2001; Peacock, 2001) and material rotations (e.g., Nur et al., 1986; Nicholson et al., 1986), shear stress is imposed across the first-generation splay fractures to form second-generation faults (right-lateral) with their associated second generation splay fractures (Fig. 9). These second-generation splay fractures, like the first-generation splay fractures, form locally between fault stepovers and outwardly from fault ends. The second generation faults almost always have a slip sense opposite to that of the first generation faults formed earlier (Fig. 9). An exception to the rule of consistent orientation and sense of slip is illustrated where an earlier splay fracture emanating from a primary left-lateral fault is more optimally oriented for the imposition of left-lateral shear (upper right, Fig. 9) and thus forms a left-lateral fault in a rather uncommon orientation. Splay fractures formed from right sense of slip across the secondgeneration faults are oriented subparallel to the first-generation left-lateral faults. This process is repeated to form successively younger generations of faults and splay fractures as slip accumulates along, and transfers between, the first generation faults (Fig. 9). The final fault network geometry is defined by the order and characteristic kink angle of splay fractures. Deviations that occur for intersection angles between faults and splay fractures might stem from the presence of complex preexisting joint geom-

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splay fracture kink angle 100 m

preexisting joint splay fracture

/

1st generation left-lateral fault 2nd generation right/left-lateral fault

3rd generation left-lateral fault 4th generation right-lateral fault

Figure 9. Conceptual model for the evolution of the strike-slip fault network at Valley of Fire. (A) Preexisting joints prior to, or at the earliest phase of, faulting. (B–E) Progressive stages of splay fractures that evolve into hierarchical sets of left- and rightlateral faults. Figure adapted from Flodin and Aydin (2004).

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etries, the formation of splay fractures at lower angles between left- and right-lateral faults, and/or material rotations.

Stop 4b—Effect of Deformation Bands on Alteration and Fluid Flow

Directions to Stop 4

This stop illustrates the interplay of depositional and structural heterogeneities on alteration and fluid flow on the outcrop scale. The banded orange-and-white alteration unit is characterized by an orange stain in the finer-grained sedimentary layers and the absence of pigment in the coarser-grained layers. This banding results in an alternating sequence of ~10–20-cm-thick orange and white layers that follow depositional bedding. In addition to this depositional control, the distribution of pigment is affected by deformation bands that act as boundaries that separate compartments with internal color gradients from orange to white. The effect of deformation bands on the distribution of the orange stain reflects the lower permeability of deformation bands relative to the more permeable surrounding sandstone. Following Antonellini and Aydin (1994) and Sigda et al. (1999), the permeability of deformation bands is typically 2–3 orders of magnitude lower compared to the adjacent sandstone. Deformation bands thus form baffles to the movement of pore fluid across otherwise more permeable sandstone. The distribution of pigment in banded orange-and-white sandstone may thus give a visual cue to the permeability structure of the sandstone, being controlled by both depositional layering and deformation bands. Assuming that the sandstone is reddened preferentially on the upstream side of deformation bands, Eichhubl et al. (2004) inferred a NE- and downward-directed flow direction for this location. A SW- and upward-directed flow direction was obtained at a location 1.5 km to the NW. They concluded that the alteration pattern provides consistent flow directions on the outcrop scale but varies on the regional scale. Walk 180 m SE to a N-S–trending gully. The east side of the gully exposes the contact of banded orange-and-white alteration at the base against yellow and purple sandstone at the top.

Follow the park road 0.6 mi (1 km) to the next parking area (no. 2) to the left of the road. Park your vehicle, cross to the east-side of the road, and climb atop the hills 180 m N of the parking area. Stop 4a—Upper Aztec Sandstone Alteration Overlook Looking north, this vantage point provides a cross-sectional view across the upper alteration units of the Aztec Sandstone (Fig. 3). Yellow and banded orange-and-white alteration colors are visible in the foreground. A layered sequence of purple, yellow, orange, white, and red colors forms the uppermost part of the section in the background. Red sandstone forms the top of White Dome to the NNW. This uniform red sandstone alteration is similar to the lower red alteration unit in the lower part of the Aztec Sandstone and is thus designated as upper red alteration. The layered alteration sequence is, at first approximation, parallel to bedding, which dips to the right (ENE). In the far background to the north, Tearfault Mesa forms a klippe of the Willow Tank thrust sheet. A closer look at the thrust and associated alteration is provided at Stop 5c. Within the regional cross section (Fig. 4), the yellow colors in the foreground are part of a roughly symmetric alteration pattern with yellow and banded orange-and-white units in the center, overlain and underlain by purple and yellow and the red sandstone of the upper and lower red units. The symmetry is broken by the occurrence of white and orange bands along the boundary to the upper red unit, and by the banded red-and-white sandstone above the upper red unit, visible between White Dome and the park road to the east. The central occurrence of yellow and purple sandstone forms a lobe that occupies most of the middle alteration units to the southeast and tapers off to the northwest. Although the alteration units are orientated roughly parallel to bedding, the top boundary of this lobe is less inclined than the other alteration bands, allowing the overlying banded orange-and-white unit to increase in thickness toward the northwest (Fig. 4). The orientation of the alteration units roughly parallel to the stratigraphic sequence suggests that the flow of fluids causing this alteration was to a large extent controlled by stratigraphic boundaries and bedding (Fig. 5). The Aztec Sandstone is overlain by mudstone of the Willow Tank Formation, which likely acted as an aquitard or seal for the permeable Aztec Sandstone aquifer. In addition, the regional scale distribution of alteration and fluid flow was affected by faults, as seen at closer range at Stop 5b. Walk east 270 m across a sand-covered area to the next exposure of banded orange-and-white sandstone. Notice deformation bands forming vertical fins on bare rock surfaces due to the higher erosion resistance of deformation bands relative to the surrounding sandstone.

Stop 4c—Crosscutting Relations of Second and Third Alteration Stages The contact between banded orange-and-white and yellow and purple alteration provides a relative age constraint on the timing of alteration and, by association, of fluid flow. The yellow and purple alteration cuts across the banded orange-and-white alteration, suggesting that the banded orange-and-white alteration predates the yellow and purple alteration. In addition, the banded orange-and-white alteration is altered to purple-and-white along this contact, also suggesting that the yellow and purple alteration postdates the banded orange-and-white alteration. We interpret this alteration pattern to be the result of dissolution and reprecipitation of iron oxide and hydroxide cement. Increasing steps of dissolution and reprecipitation lead to a coarsening of iron oxide and hydroxide, with a resulting change in color from red and orange to yellow and purple. We attribute the banded orange-andwhite stain to a second stage of alteration, distinct from a third stage of alteration that resulted in yellow and purple alteration.

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The three alteration stages at Valley of Fire therefore include (1) a syndepositional stage resulting in the uniform red stain of the lower and upper red alteration units; (2) a second stage of alteration resulting in banded orange-and-white alteration; and (3) a third stage characterized by yellow and purple colors (Fig. 5). Walk 50 m to a deep N-S–trending canyon. This canyon follows one branch of the Lonewolf fault we examined farther south at Stops 3a–3c. A second branch of the fault is in the parallel canyon 20 m east.

the purple and, by association, yellow alteration largely postdates faulting along these high-angle faults of inferred Tertiary age.

Stop 4d—Segmented Faults and Meso-Scale Material Rotation

Stop 5a—Preferred Fluid Flow and Alteration along Sheared Joints

The Lonewolf fault offsets the lower contact of the upper red sandstone by 33 m of left-lateral slip. Similar to the upper contact of the lower red sandstone, this alteration contact predates fault slip along these high-angle faults. Here, we see another example of material rotation between segmented strike-slip faults, though larger in scale and opposite in rotation sense compared to Stop 3b. In this case, the bounding faults are left-lateral and the internal faults are right-lateral. The bounding faults are the right-stepping northern and central segments of the Lonewolf fault. The internal right-lateral faults intersect their bounding faults at angles ranging from 27° to 60°. Oriented perpendicular to, and offset by, the two segments of the left-lateral Lonewolf fault is a zone of earlier-formed deformation bands. The zone of deformation bands itself consists of bounding deformation bands that show right-lateral separation and internal deformation bands that show left-lateral separation and are similar in geometry to zones of deformation bands described by Davis et al. (2000). The mean angle of intersection between the two deformation band sets is ~26° both inside and outside of the bounding segments of the Lonewolf fault. However, mean orientations of the two sets that occur between the fault segments are rotated ~15° in a counterclockwise sense with respect to the zones of deformation bands found outside of these segments. Some of the internal right-lateral faults appear to be rotated as well. Walk 180 m south along the trace of the Lonewolf fault, to the south end of the canyon.

Following the road back from parking area 3, we crossed from upper red Aztec Sandstone through some white alteration into a 10–20-m-thick band of orange sandstone (Fig. 3). We also note that the alteration colors do not appear to continue across the road with orange sandstone on the west side against yellow and purple sandstone on the east side. This mismatch is due to slip along the N-S–striking Bighorn fault. This fault, which follows the road from the last turn onward, offsets the orange, white, and red alteration units by 600–700 m of apparent left-lateral strikeslip. A purple band ~30 m south of the yellow to orange contact also appears offset along the fault but only by 10–20 m. An exact amount is difficult to determine due to the presence of the road. In accordance with our finding at the Lonewolf fault, we can conclude, however, that the yellow and purple alteration formed later than the orange, white, and upper red alteration units—the former during late stages of faulting, the latter prior to faulting. Joints across the lower contact of the orange alteration zone and extending into the underlying white sandstone are characterized by orange alteration halos. In some instances, the lower boundary of the orange alteration zone is deflected downward along these joints as elongate orange lobes, consistent with these joints acting as preferred conduits for fluid flow and mass transfer. The preferred staining along these joints also indicates that the orange stain is secondary to the bleaching of the white sandstone. Following the observation that the orange alteration unit is offset by the same amount as the white and upper red contacts, and thus predated fault slip, we interpret the orange halos around these joints as a secondary remobilization of the orange stain after faulting initiated. These alteration halos are in part offset by 1–3 cm of left-lateral slip, kinematically compatible with slip along the Bighorn fault and thus interpreted as coeval with fault movement. Follow the Bighorn fault back north, up the hill along the road. Where the road swings northwest, leave the road and continue straight along the fault for 160 m.

Stop 4e—A Fault Core with ~25 m Slip and Low Average Shear Strain At this stop, we will visit another outcrop of the Lonewolf fault core. As at Stop 3a, the Lonewolf fault at this locality has ~25 m of left-lateral slip. However, the fault core here is much wider than the core at Stop 3a, resulting in a lower value of average shear strain across the zone. The lower shear strain appears to have an impact on fault rock petrophysics. Of the three sampling stations, this locality exhibits the least amount of permeability reduction. Unlike the lower contact of the upper red sandstone, which is offset 33 m along this fault, purple bands at this location cross the fault with minimal, if any, mechanical offset. We conclude that

Directions to Stop 5 Return to the vehicles. Follow the park road north 1.1 mi (1.8 km) and park at parking area 3. Walk back along the road 830 m; Stop 5a is 10 m to the west of, and below, the road, ~130 m north of Kaolin Wash.

Stop 5b—Influence of Bighorn Fault on Alteration and Fluid Flow Following the Bighorn fault northward, the fault separates red and banded red-and-white sandstone to the west (left) from

Brittle deformation, fluid flow, and diagenesis in sandstone yellow and purple sandstone to the east. Walking up the hill, we observe purple and orange alteration bands on the east side of the fault changing orientation adjacent to the fault, turning from an orientation roughly parallel to bedding into parallelism upon approaching the fault. This indicates that the reactions leading to the formation of these bands were affected by the presence of the Bighorn fault, thus dating these reactions to after, or concurrent with, the formation of the fault. Unlike the mechanical fault offset observed for contacts of the lower and upper red sandstone and for banded orange-and-white sandstone, this deflection of the purple and orange bands is not due to shearing along the fault. Instead, we interpret this deflection to be the result of remobilization and “bleeding” of iron across the fault where the fault juxtaposes red sandstone against bleached sandstone. A similar deflection of purple and yellow bands is observed along faults south of Kaolin Wash and along the Classic fault. Continue along the Bighorn fault. Upon approaching a gravel-capped low hill to the left, turn east for ~300 m to obtain a good view of Tearfault Mesa to the north and of the alteration of the underlying Aztec Formation. Stop 5c—Tearfault Mesa Overlook: Effect of Thrusting on Alteration and Fluid Flow Tearfault Mesa is a tectonic klippe of allochthonous Aztec Sandstone, resting on gravel, sandstone, and gray mudstone of the Cretaceous Willow Tank Formation. The Willow Tank Formation, in turn, overlies in depositional contact the Aztec Sandstone that forms the bulk of Aztec exposure in the park (Figs. 3 and 4). This klippe is part of the Willow Tank thrust sheet, the lowermost and easternmost Sevier thrust sheet in this area (Longwell, 1949; Bohannon, 1983). Along the eastern tip of Tearfault Mesa, the beds of the Willow Tank Formation are locally overturned, reflecting the eastward movement of the thrust front. An E-W–striking tear fault, which is difficult to discern from this vantage point, forms the southern boundary of the klippe. The Bighorn fault cuts across and offsets vertically the Willow Tank thrust sheet, consistent with a Tertiary age of the Bighorn fault. The upper red alteration of the Aztec Sandstone is bleached white in the vicinity of the thrust. This suggests that (1) the uniform red color of the upper red sandstone, considered the first alteration stage, predates the thrusting; and (2) that bleaching of the upper red is synchronous to or postdates thrusting. A pre-thrust age of red alteration of the Aztec Sandstone was also deduced by Longwell (1949) based on the ubiquitous occurrence of red Aztec clasts in the Willow Tank Formation. Although no direct cross-cutting relations are observed between the bleaching on top of the upper red alteration unit with the other alteration units, we correlate this bleaching with the fluid flow event that caused the banded orange-and-white alteration at stratigraphically and structurally deeper levels. These alteration stages are mechanically offset by the high-angle faults of Tertiary age, thus predating these faults.

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Directions to Stop 6 Return to vehicles and follow the park road to its end at the White Dome parking area. Walk back along the road for 200 m, then turn left (W), following a marked trail for 100 m. Stop 6a—Deformation Bands as Flow Barriers The south- to west-facing rock exposure contains the lower contact of the orange alteration unit observed farther east at Stop 5a. At this location, this alteration contact appears offset along deformation bands by 15–20 cm. Taylor and Pollard (2000) noted that these offsets are not due to slip along these bands, which rarely exceeds a few millimeters. Rather, these offsets are due to the retardation of a reaction boundary across these bands relative to the undeformed sandstone. Taylor and Pollard (2000) attributed this retardation to the reduced permeability of these bands. Assuming that these bands would retard the velocity of a sweeping reaction boundary to the same degree as the fluid flow velocity, they calculated a reduction in permeability across these bands by one to two orders of magnitude. This estimate is approximately one to two orders of magnitude lower than measurements by Antonellini and Aydin (1994) for deformation bands in similar eolian sandstone. We note that the velocity of a moving reaction boundary would, in fact, be slower than the flow velocity of the reacting aqueous solution. The calculated permeability reduction is thus a minimum estimate of the actual permeability reduction, which could account for the difference from the measured values by Antonellini and Aydin (1994). Return to the parking area and walk along the marked hiking trail to the south of the parking area through the wind gap west of White Dome. Where the trail drops into the drainage to the south, leave the trail and follow the sandstone ledges to your left, gaining some height, to the ridge south of White Dome. Use caution walking on the slick rock surfaces! Stop 6b—Summary Overlook of Valley of Fire Alteration and Fluid Flow This vantage point provides a sweeping view over the northern part of Valley of Fire, with the tectonic klippe of Tearfault Mesa to the east and Silica Dome, marking the contact of lower red to middle alteration units, to the far southeast. We are standing on yellow and purple alteration that cuts across banded orange-and-white alteration. Yellow and purple alteration is overlain by orange and white alteration units, and, close to the top of White Dome, a bit of upper red. The alteration units are overall parallel or subparallel to bedding, which dips NE at 20°. This view summarizes well the complex interaction of diagenetic alteration and fluid flow with depositional and structural heterogeneities in the Aztec Sandstone at Valley of Fire: On the outcrop scale, alteration and fluid flow are influenced and controlled by bedding, deformation bands, and joints. On the regional scale, alteration and fluid flow are controlled by

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formation contacts and large scale structures, including the Willow Tank thrust, with Tearfault Mesa as its erosional remnant. Although the depositional, structural, hydrogeologic, and diagenetic history is challenging to decipher, Valley of Fire provides a perhaps unique, and certainly picturesque, view into the complex interaction among these processes in the subsurface. Return along the trail to the White Dome parking area. Return to Las Vegas. ACKNOWLEDGMENTS We thank A. Aydin and D.D. Pollard, Stanford University, for introducing us to Valley of Fire geology, promoting this research, and for sharing their ideas. Material presented here includes work performed by former Stanford students and postdocs including J. Chapin, H. Jourde, R. Myers, K. Sternlof, and W.L. Taylor. This research was funded by U.S. Department of Energy grant DE-FG 03-94ER14462, with additional funding from the Stanford Rock Fracture Project (RFP) industrial affiliate program, Chevron Energy Technology Company, and Stanford University. REFERENCES CITED Anderson, R.E., and Barnhard, T.P., 1993, Heterogeneous Neogene strain and its bearing on horizontal extension and horizontal and vertical contraction at the margin of the extensional orogen, Mormon Mountains area, Nevada and Utah: U.S. Geological Survey Bulletin, B2011, 5 sheets, 113 p. Antonellini, M.A., and Aydin, A., 1994, Effects of faulting on fluid flow in porous sandstones: petrophysical properties: AAPG Bulletin, v. 78, p. 355–377. Axen, G.J., 1993, Ramp-flat detachment faulting and low-angle normal reactivation of the Tule Springs thrust, southern Nevada: Geological Society of America Bulletin, v. 105, p. 1076–1090, doi: 10.1130/00167606(1993)105<1076:RFDFAL>2.3.CO;2. Aydin, A., 2000, Fractures, faults, and hydrocarbon migration and flow: Marine and Petroleum Geology, v. 17, p. 797–814, doi: 10.1016/S02648172(00)00020-9. Blakey, R.C., 1989, Triassic and Jurassic geology of the southern Colorado Plateau, in Jenny, J.P., and Reynolds, S.J., eds., Geologic evolution of Arizona: Arizona Geological Society Digest, v. 17, p. 369–396. Bohannon, R.G., 1977, Geologic map and sections of the Valley of Fire region, North Muddy Mountains, Clark County, Nevada: U.S. Geological Survey Miscellaneous Field Studies Map MF-849, scale 1:25,000. Bohannon, R.G., 1983, Mesozoic and Cenozoic tectonic development of the Muddy, North Muddy, and northern Black Mountains, Clark County, Nevada: Geological Society of America Memoir 157, p. 125–148. Bohannon, R.G., 1992, Geologic map of the Weiser Ridge quadrangle, Clark County, Nevada: U.S. Geological Survey Report GQ-1714, 1 sheet, scale 1:24,000. Bohannon, R.G., Grow, J.A., Miller, J.J., and Blank, R.H., Jr., 1993, Seismic stratigraphy and tectonic development of Virgin River depression and associated basins, southeastern Nevada and northwestern Arizona: Geological Society of America Bulletin, v. 105, p. 501–520, doi: 10.1130/ 0016-7606(1993)105<0501:SSATDO>2.3.CO;2. Bourne, S.J., and Willemse, E.J.M., 2001, Elastic stress control on the pattern of tensile fracturing around a small fault network at Nash Point, UK: Journal of Structural Geology, v. 23, p. 1753–1770, doi: 10.1016/S01918141(01)00027-X. Caine, J.S., Evans, J.P., and Forster, C.B., 1996, Fault zone architecture and permeability structure: Geology, v. 24, p. 1025–1028, doi: 10.1130/00917613(1996)024<1025:FZAAPS>2.3.CO;2. Carpenter, D.G., 1989, Geology of the North Muddy Mountains, Clark County Nevada, and regional structural synthesis: fold-thrust and Basin-Range

structure in southern Nevada, southwest Utah and northwest Arizona [M.S. thesis]: Corvallis, Oregon State University, 145 p. Carpenter, D.G., and Carpenter, J.A., 1994, Fold-thrust structure, synorogenic rocks, and structural analysis of the North Muddy and Muddy Mountains, Clark County, Nevada, in Dobbs, S.W., and Taylor, W.J., eds., Structural and stratigraphic investigations and petroleum potential of Nevada, with special emphasis south of the Railroad Valley Producing Trend: Nevada Petroleum Society 1994 Conference Volume II: Reno, Nevada, Nevada Petroleum Society, p. 65–94. Davis, G.H., Bump, A.P., Garcia, P.E., and Ahlgren, S.G., 2000, Conjugate Riedel deformation band shear zones: Journal of Structural Geology, v. 22, p. 169–190, doi: 10.1016/S0191-8141(99)00140-6. Dickinson, W.W., and Milliken, K.L., 1995, The diagenetic role of brittle deformation in compaction and pressure solution, Etjo Sandstone, Namibia: Journal of Geology, v. 103, p. 339–347. Eichhubl, P., and Behl, R.J., 1998, Diagenesis, deformation, and fluid flow in the Miocene Monterey Formation of coastal California, in Eichhubl, P., ed., Diagenesis, deformation, and fluid flow in the Miocene Monterey Formation of coastal California: Society for Sedimentary Geology (SEPM) Pacific Section Special Publication 83, p. 5–13. Eichhubl, P., and Boles, J.R., 2000, Focused fluid flow along faults in the Monterey Formation, coastal California: Geological Society of America Bulletin, v. 112, p. 1667–1679, doi: 10.1130/0016-7606(2000)112<1667: FFFAFI>2.0.CO;2. Eichhubl, P., Taylor, W.L., Pollard, D.D., and Aydin, A., 2004, Paleofluid flow and deformation in the Aztec Sandstone at the Valley of Fire, Nevada— Evidence for the coupling of hydrogeologic, diagenetic, and tectonic processes: Geological Society of America Bulletin, v. 116, p. 1120–1136, doi: 10.1130/B25446.1. Flodin, E.A., and Aydin, A., 2004, Evolution of a strike-slip fault network, Valley of Fire State Park, southern Nevada: Geological Society of America Bulletin, v. 116, p. 42–59, doi: 10.1130/B25282.1. Flodin, E.A., Aydin, A., Durlofsky, L.J., and Yeten, B., 2001, Representation of fault zone permeability in reservoir flow models: SPE paper 71671, presented at the Society of Petroleum Engineers Annual Technical Conference and Exhibition, New Orleans, 10 p. Flodin, E.A., Prasad, M., and Aydin, A., 2003, Petrophysical constraints on deformation styles in Aztec Sandstone: Pure and Applied Geophysics, v. 160, p. 1589–1610. Flodin, E.A., Gerdes, M., Aydin, A., and Wiggins, W.D., 2005, Petrophysical properties and sealing capacity of fault rock, Aztec Sandstone, Nevada, in Sorkhabi, R., and Tsuji, Y., eds., Faults, fluid flow, and petroleum traps: American Association for Petroleum Geologists Memoir, v. 85, p. 197–217. Gale, J.F.W., Laubach, S.E., Marrett, R.A., Olson, J.E., Holder, J., and Reed, R.M., 2004, Predicting and characterizing fractures in dolostone reservoirs: Using the link between diagenesis and fracturing, in Braithwaite, C.J.R., Rizzi, G., and Darke, G., eds., The geometry and petrogenesis of dolomite hydrocarbon reservoirs: London, Geological Society Special Publication 235, p. 177–192. Gross, M.R., 1995, Fracture partitioning: Failure mode as a function of lithology in the Monterey Formation of coastal California: Geological Society of America Bulletin, v. 107, p. 779–792, doi: 10.1130/00167606(1995)107<0779:FPFMAA>2.3.CO;2. Hill, R.E., 1989, Analysis of deformation bands in the Aztec Sandstone, Valley of Fire State Park, Nevada [M.S. thesis]: Las Vegas, University of Nevada, 68 p. Jourde, H., Flodin, E.A., Aydin, A., Durlofsky, L.J., and Wen, X.-H., 2002, Computing permeability of fault zones in aeolian sandstone from outcrop measurements: AAPG Bulletin, v. 86, p. 1187–1200. Kattenhorn, S.A., Aydin, A., and Pollard, D.D., 2000, Joints at high angles to normal fault strike: An explanation using 3-D numerical models of faultperturbed stress fields: Journal of Structural Geology, v. 22, p. 1–23, doi: 10.1016/S0191-8141(99)00130-3. Laubach, S.E., Reed, R.M., Olson, J.E., Lander, R.H., and Bonnell, L.M., 2004, Coevolution of crack-seal texture and fracture porosity in sedimentary rocks: Cathodoluminescence observations of regional fractures: Journal of Structural Geology, v. 26, p. 967–982, doi: 10.1016/j.jsg.2003.08.019. Lin, P., and Logan, J.M., 1991, The interaction of two closely spaced cracks: A rock model study: Journal of Geophysical Research, v. 96, p. 21,667–21,675. Longwell, C.R., 1949, Structure of the northern Muddy Mountain area, Nevada: Geological Society of America Bulletin, v. 60, p. 923–967.

Brittle deformation, fluid flow, and diagenesis in sandstone Martel, S.J., 1990, Formation of compound strike-slip fault zones, Mount Abbot quadrangle, California: Journal of Structural Geology, v. 12, p. 869–882, doi: 10.1016/0191-8141(90)90060-C. Marzolf, J., 1983, Changing wind and hydraulic regimes during deposition of the Navajo and Aztec sandstones, Jurassic (?) southwestern United States, in Brookfield, M.E., and Ahlbrandt, T.S., eds., Eolian sediments and processes: Amsterdam, Elsevier, p. 635–660. Marzolf, J.E., 1990, Reconstruction of extensionally dismembered early Mesozoic sedimentary basins; Southwestern Colorado Plateau to the eastern Mojave Desert, in Wernicke, B.P., ed., Basin and Range extensional tectonics near the latitude of Las Vegas, Nevada: Geological Society of America Memoir 176, p. 477–500. Mollema, P.N., and Antonellini, M.A., 1996, Compaction bands: a structural analog for anti-mode I cracks in aeolian sandstone: Tectonophysics, v. 267, p. 209–228, doi: 10.1016/S0040-1951(96)00098-4. Munsell Soil Color Charts, 1994, revised edition: New Windsor, New York, Munsell Color, 10 p. Myers, R.D., 1999, Structure and hydraulics of brittle faults in sandstone [Ph.D. thesis]: Stanford, Stanford University, 176 p. Myers, R.D., and Aydin, A., 2004, The evolution of faults formed by shearing across joint zones in sandstone: Journal of Structural Geology, v. 26, p. 947–966, doi: 10.1016/j.jsg.2003.07.008. Nicholson, C., Seeber, L., Williams, P., and Sykes, L.R., 1986, Seismic evidence for conjugate slip and block rotation within the San Andreas fault system, southern California: Tectonics, v. 5, p. 629–648. Nur, A., Ron, H., and Scotti, O., 1986, Fault mechanics and the kinematics of block rotations: Geology, v. 14, p. 746–749, doi: 10.1130/00917613(1986)14<746:FMATKO>2.0.CO;2. Peacock, D.C.P., 2001, The temporal relationship between joints and faults: Journal of Structural Geology, v. 23, p. 329–341, doi: 10.1016/S01918141(00)00099-7. Poole, F.G., 1964, Paleowinds in the western United States, in Nairn, A.E.M., ed., Problems in paleoclimatology: London, John Wiley, p. 394–406.

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Rock-Color Chart Committee, 1970, Rock-color chart, reprint: Boulder, Colorado, Geological Society of America, 16 p. Segall, P., and Pollard, D.D., 1980, Mechanics of discontinuous faults: Journal of Geophysical Research, v. 85, p. 4337–4350. Sigda, J.M., Goodwin, L.B., Mozley, P.S., and Wilson, J.L., 1999, Permeability alteration in small-displacement faults in poorly lithified sediments: Rio Grande Rift, Central New Mexico, in Haneberg, W.C., Mozley, P.S., Moore, J.C., and Goodwin, L.B., eds., Faults and Subsurface fluid flow in the shallow crust: Geophysical Monograph, v. 113, Washington D.C., American Geophysical Union, p. 51–68. Sternlof, K., Chapin, J.R., Pollard, D.D., and Durlofsky, L.J., 2004, Permeability effects of deformation band arrays in sandstone: AAPG Bulletin, v. 88, p. 1315–1329, doi: 10.1306/03280403103. Sternlof, K., and Pollard, D.D., 2001, Deformation bands as linear elastic fractures: progress in theory and observation: Eos (Transactions, American Geophysical Union), v. 82, no. 47, Fall Meeting Supplement, Abstract T42E-04. Taylor, W.L., 1999, Fluid flow and chemical alteration in fractured sandstones [Ph.D. thesis]: Stanford, California, Stanford University, 411 p. Taylor, W.L., and Pollard, D.D., 2000, Estimation of in situ permeability of deformation bands in porous sandstone, Valley of Fire, Nevada: Water Resources Research, v. 36, p. 2595–2606, doi: 10.1029/2000WR900120. Taylor, W.L., Pollard, D.D., and Aydin, A., 1999, Fluid flow in discrete joint sets: Field observations and numerical simulations: Journal of Geophysical Research, v. 104, p. 28983–29006, doi: 10.1029/1999JB900179. Worden, R.H., and Morad, S., 2000, Quartz cementation in oil field sandstones: a review of key controversies, in Worden, R.H., and Morad, S., eds., Quartz cementation in sandstones: International Association of Sedimentologists Special Publication no. 29: Oxford, UK, Blackwell Science, p. 1–20. Worden, R.H., and Morad, S., 2003, Clay minerals in sandstones: controls on formation, distribution and evolution, in Worden, R.H., and Morad, S., eds., Clay mineral cements in sandstones: International Association of Sedimentologists Special Publication no. 34: Oxford, UK, Blackwell Science, p. 3–41.

Printed in the USA

Geological Society of America Field Guide 6 2005

Evolution of a late Cenozoic supradetachment basin above a flat-on-flat detachment with a folded lateral ramp, SE Idaho Alexander N. Steely Department of Geology, Utah State University, Logan, Utah 84322-4505, USA Susanne U. Janecke Department of Geology, Utah State University, Logan, Utah 84322-4505, USA Sean P. Long Department of Geosciences, University of Idaho, Moscow, Idaho 83844, USA Stephanie M. Carney Department of Geology, Utah State University, Logan, Utah 84322-4505, USA Robert Q. Oaks Jr. Department of Geology, Utah State University, Logan, Utah 84322-4505, USA Victoria E. Langenheim U.S. Geological Survey, 345 Middlefield Road, Menlo Park, California 94025, USA Paul K. Link Department of Geosciences, Idaho State University, P.O. Box 8072, Pocatello, Idaho 83209, USA

ABSTRACT Uplift and exposure of the Bannock detachment system and the synextensional basin fill deposits of the Salt Lake Formation provide a unique exposure of the threedimensional geometries of a low-angle normal fault system and the stratal architecture of the overlying supradetachment basin. Within this system, structural and stratigraphic analyses, outcrop patterns, tephra geochronology, and geological cross sections document several important relationships: (1) the Bannock detachment system developed above the Sevier-age Cache-Pocatello culmination and resembles the Sevier Desert detachment in its geometry, structural setting, and kinematic evolution; (2) the Bannock detachment system initiated and slipped at low angles; (3) flaton-flat, ramp-flat, and lateral ramp geometries, as well as excision, can significantly affect the hanging wall deformation style due to the shallow depth (~2–4 km) of the Bannock detachment fault during late stages of slip; (4) late Miocene–Pliocene tuffaceous synrift deposits of the Salt Lake Formation record deposition in a supradetachment basin, display an unroofing sequence, and a three-stage evolution that includes pre-translation, translation, and breakup phases. Recycled pre-translation and transSteely, A.N., Janecke, S.U., Long, S.P., Carney, S.M., Oaks Jr., R.Q., Langenheim, V.E., and Link, P.K., 2005, Evolution of a late Cenozoic supradetachment basin above a flat-on-flat detachment with a folded lateral ramp, SE Idaho, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 169–198, doi: 10.1130/2005.fld006(08). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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lation phase deposits are diagnostic of this evolution; and (5) beginning in mid- to late Pliocene time, high-angle, north-striking Basin and Range faults disrupted and dismembered the Bannock detachment system. Keywords: detachment, basin evolution, lateral ramp, Salt Lake Formation, extensional fold.

INTRODUCTION

On this field trip, we will visit critical exposures of the late Miocene–Pliocene Bannock detachment system and its associated supradetachment basin in southeastern Idaho (Fig. 1). We illustrate several field relationships in the Malad City East and Clifton quadrangles that were described previously in Janecke and Evans (1999), Janecke et al. (2003), Carney and Janecke (2005), and Carney et al. (2004), and these sources provide more detailed information. Our main focus is on two more recent studies in the Weston Canyon and Henderson Creek quadrangles that test, update, and refine earlier analyses. New geologic mapping in the Weston Canyon, Henderson Creek, Clifton, and Malad City East quadrangles forms the core of our data sets (Carney et al., 2004; Long et al., 2004; Steely and Janecke, 2005; Evans et al., 2000). Hanging wall and footwall exposures of the Clifton strand of the Bannock detachment system along Rattlesnake Ridge in the Weston Canyon quadrangle comprise the majority of stops on this trip. Initial work provided evidence for slip on a subhorizontal normal fault in the form of a flat-on-flat low-angle detachment (Carney and Janecke, 2005), and subsequent mapping and analysis refine this relationship. At Stops 5 through 9, we examine this now well established relationship, discuss its implications, and investigate a ~850–1100-m-high folded lateral ramp along the same detachment fault. We also examine basin-fill deposits that record early translation and later breakup of the detachment’s hanging wall. A secondary focus is new work in the Henderson Creek quadrangle. Field studies in the southwest edge of the late Miocene–Pliocene Salt Lake Formation sedimentary basin identified and dated syntectonic conglomerates shed from the antithetic Steel Canyon normal fault in the hanging wall of the Bannock detachment fault (Long, 2004). Stop 1 briefly describes some of the key results of Long’s research (Long, 2004; Long et al., 2004).

is NNW-trending, 150 km long, and 25–30 km wide before Cenozoic extension. It exposed lower Cambrian rocks in its core (Fig. 2) and was first recognized by Rodgers and Janecke (1992). The culmination was named and refined by Carney (2002) and Long (2004). It likely formed east of the frontal Malad ramp of the Paris-Willard thrust system as a hanging wall ramp anticline on a footwall flat. At least two phases of Cenozoic extension followed the Sevier Orogeny in the western United States (Stewart, 1998). Typically, the first phase of extension was accommodated by listric- to low-angle normal faults that developed in the hinterland and fold-and-thrust belt of the Cordillera. This extensional episode produced Eocene-Miocene metamorphic core complexes such as the top-to-the-east Raft River-Albion-Grouse Creek metamorphic core complex located ~125 km west of the Bannock detachment system (Fig. 3) (Wells et al., 2000). The structural style of the Bannock detachment system resembles this early phase of extension but the age and the magnitude of extension is less than that of most core complexes (Carney and Janecke, 2005). The second major phase of extension in the western United States began in late early Miocene to Pliocene time and was accommodated by moderately- to steeply-dipping, planar to listric normal faults. These faults bound present-day mountain ranges that trend ~N-S (Figs. 1 and 4) and are regularly spaced ~30 km apart between the Sierra Nevada ange and the Colorado Plateau. Recent syntheses suggest that initiation of this phase of Basin and Range extension was not synchronous, but rather has migrated both east and west through time (Stewart, 1998; Henry and Perkins, 2001; Stockli, 2000; Janecke et al., 2003). The Bannock and Malad Ranges are both uplifted on the east and west by active north-striking Basin and Range normal faults. The Malad Range is uplifted on the west by northern segments of the Wasatch Fault, which extends southward 370 km through Utah (Machette et al., 1992; Chang and Smith, 2002). A new gravity data set presented here shows that the two normal fault zones that bound Cache Valley, the next basin to the east, have throws similar to the northern Wasatch fault (Oaks et al., 2005).

Regional Geology

Tectonic Problems along Low-Angle Normal Faults

The late Cenozoic Bannock detachment system developed at the transition between the preexisting Cordilleran hinterland and the fold-and-thrust belt. A pre-Tertiary subcrop map of this area shows that the detachment formed above a major culmination related to uplift above the Paris-Willard thrust system during the Sevier Orogeny (Fig. 2). The Cache-Pocatello culmination

Although low-angle normal fault systems have been widely recognized throughout the Basin and Range province, they are not uniformly distributed (e.g., Wernicke, 1992; Stewart, 1998). Some have suggested that they localize near “contractional welts” or culminations in the lithosphere (Coney and Harms, 1984; Spencer and Reynolds, 1991). The Bannock detachment system

Field Trip Goals

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Figure 1. Simplified geologic map showing the distribution of the Salt Lake Formation around Cache Valley and the active Basin and Range normal faults. Some older normal faults are also shown (gray). The Clifton horst is the up-thrown block between the Deep Creek and Dayton-Oxford faults. CV—Cottonwood Valley; DCF—Deep Creek Fault; DCHG— Deep Creek half graben; DOF—Dayton-Oxford Fault; ECF—East Cache Fault (N—northern; C—central; S—southern segment); OP—Oxford Peak; Q—Quaternary sediment; RRP—Red Rock Pass; Tsl—Salt Lake Formation; WCF—West Cache fault zone (CM—Clarkston Mountain; JH—Junction Hills; W—Wellsville segment); WF—Wasatch Fault. Some buried faults (dotted) from Zoback (1983). This map is modified from Janecke et al. (2003), with updates from unpublished gravity data of Oaks et al., 2005 (see Fig. 15) and Eversaul (2004).

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Marsh Valley

Ma

has a close spatial association with the Cache-Pocatello culmination (Fig. 3), and Carney and Janecke (2005) proposed that the additional gravitational potential energy and residual topography associated with the culmination helped to localize the Bannock detachment system. The collapse of culminations is a recurrent theme in extensional orogens that deform older contractional belts, and there are many examples of major normal faults localized above culminations within the western United States including the Sevier Desert detachment, Wasatch fault, and Salmon, Idaho, area detachment faults (Fig. 3) (Coogan and DeCelles, 1996; Yonkee, 1992; Janecke et al., 2000; Janecke and Blankenau, 2003). The Bannock and Sevier Desert detachment systems are similar in structural style; both are located 100–125 km east of metamorphic core complexes, and both developed above regional Sevier-age culminations (Fig. 3) (Carney and Janecke, 2005, and references therein). Three models may explain the geological record of lowangle normal faults (Fig. 5A). Such faults may form and slip at low angles (Wernicke, 1981; Allmendinger et al., 1983; Lister and Davis, 1989; Davis and Lister, 1988), rotate to lower angles via domino-style extension (Proffett, 1977; Wernicke and Burchfiel, 1982; Gans and Miller, 1983; Rodgers et al., 2002), or back-rotate isostatically at a rolling-hinge (Buck, 1988; Wernicke and Axen, 1988; Brady et al., 2000). A specific suite of structures and relationships is predicted if a detachment fault initiates and slips at low angles (Wernicke, 1985; Davis and Lister, 1988; Carney and Janecke, 2005). Some hanging walls of detachment faults are structurally intact, but, as the amount of extension increases, some hanging walls break up internally along planar and listric high- to low-angle normal faults that are either cut by, or sole into, the master fault (Wernicke, 1985, 1992; Fedo and Miller, 1992; Fowler et al., 1995; Janecke et al., 2003; Carney and Janecke, 2005). As overburden thickness decreases in the hanging wall due to crustal extension, footwall mid-crustal rocks, as well as the bounding master fault, may dome upward to create a broad antiform at high angles to the extension direction (Spencer, 1984). During this process, the master fault may cut up into (excise) hanging wall rocks and may produce a secondary breakaway (Fig. 5B) (John, 1987; Davis and Lister, 1988; Carney and Janecke, 2005). Excision

171

5

10

15

15 20

mi

km

Quaternary-Tertiary deposits undifferentiated

Wasatch Mts. town 91

road

Tertiary Salt Lake Formation pre-Tertiary bedrock Neoproterozoic Pocatello Formation

approximate attitude of bedding stratigraphic contact

active normal fault

low-angle normal fault

concealed normal fault

older inactive normal fault

112˚ 00’ A.N. Steely et al.

172 45’ 42˚

42˚ 45’

I-15

control point on Tertiary unconformity

ust thr

st hru

Range

is Par

t nam Put

Bannock

Gem ver r Ri a e B

r Po

Modern mountain ranges

tn f eu

34

91

Valley

Marsh

Road Anticline

Ra

Syncline

e

fau

ng

Valley

y lle Va

91

lt

Overturned syncline

Ban noc

k

fault

Bear River

ne

C-TR

Loga n s.

CZb

Mt

Cbl to Cbo

syncli

91

WF

Bear River Valley

lPzu

East

le

Cn to Csc

Bear Lake Range

il sv

Ordovician rocks

Bear River

ll We

I-84

42˚ 00’

Peak

culmination

mp

Mississippian to Permian rocks Silurian to Devonian rocks

West Hills

a ad R

Mal

Triassic rocks

0

41˚ 30’ Lower Paleozoic rocks undivided (lPzu) Outline of CachePocatello culmination

t

Valley

I-15

thrus

Idaho Utah

e clin syn

Cache

er Riv

o Cache Pocatell

Malad Valley Malad Range Samaria Mountains Cl ar 42˚ 00’ ks to M n tn .

Paris

WF

amp

r B ea

en Fish Hav

ge

OPA

EDOFDOF

ad R Mal

Ran

Cache

Pleasant View Hills

112˚ 00’

10

20 km

Sevier-aged thrust fault

Tertiary normal fault

Sevier-aged contact

Pre-Tertiary normal fault

Figure 2. Subcrop map showing the location and age of rocks beneath the Tertiary unconformity. Also shown are Mesozoic folds and thrusts of the Servier Orogeny, the location of the Oxford Peak anticline (OPA), and the location of the Cache Pocatello culmination. Modified from Carney (2002). In mountain ranges, the age of the youngest exposed rock unit was also used to construct the subcrop map. The E-W extent of the culmination is 50%–60% greater than it was at the end of the Sevier thrusting because of subsequent extension. DOF—Dayton-Oxford Fault; EDOF—East Dayton-Oxford Fault; WF—Wasatch Fault.

Evolution of a late Cenozoic supradetachment basin

117 o

173

111o

WA 46

MT

o

?

ID

?

SAC

45

o

YH

OR

WSRP CPC

WY

Raft River MCC

42 o

NV Area of Fig. 1

0 0

41o

WC

100 miles

UT

100 km

Detachment fault

Snake Range MCC

Buried W-dipping thrust ramp

WF

SC

SDD

Metamorphic core complex (MCC) Western Snake River Plain Eastern Snake River Plain

Structural culminations of fold-and-thrust belt (Cambrian and older subcrop) Structural culminations of fold-and-thrust belt (Devonian and older subcrop) Areas outside the Basin and Range province

Basin-and-Range fault (sensu strictu)

Inferred lateral extent of the breakway for the Bannock detachment and Sevier Desert systems

Figure 3. Regional map showing the major active normal faults adjacent to the Eastern Snake River Plain (ESRP), the NNW-trending Cache-Pocatello culmination (CPC), and other culminations of the thrust belt. Note that the Bannock detachment system coincides with the Cache-Pocatello culmination and has the opposite vergence as the partly coeval Raft River detachment (Wells et al., 2000). A Cambrian and older pre-Tertiary subcrop defines the extent of the Sevier, Wasatch, and Salmon area and most uplifted portions of the Cache-Pocatello culminations. The extent of the Devonian and older subcrop is shown for the CPC. SC—Sevier Desert culmination; SAC—Salmon area culminations; WC—Wasatch culmination; SDD—Sevier Desert detachment; YH—Yellowstone hotspot. Modified from Carney and Janecke (2005).

Clifton quadrangle C-Zb

Tcv

C-Zb C-Zb

O-D O-D

QTg

Tcv

Tcv

QTg

O-D Q Ttc Tcv

Tcv

Clifton fault

Tnc

Tcv

O-D

Tcv Q

Ttc Q

Tcv

INT

C-O

ER

2

C-O

ILE

Ts

3

Tcv

RO AD

Q QTg QTg

Tcv

C-O

Tcv Tcv

C-O

Ts C-O

Tcv

C-O

C-O

Tcv

C-O

C-O

Malad City East quadrangle Henderson Creek quadrangle

C-O

C-O

Kno ll fa ult

C-O

Tcv

C-O

C-Zb

Zp

Zp

C-O C-O

Q

6

Tcv

Q

Tcv

Ttc-cv

C-O

C-O

C-O

ee

Tufa-bea ring Micrite-be aring

lC

an yo C-O n

4000’

0‘

fa u

Tcv

lt

0 km

1 km

Tcv

DRY CAN YON HOR ST 8000’

2 km Tscu

N

undivided Quaternary deposits

Zp

Tnc

New Canyon Member Third Creek Member and Tscu Upper conglomerate unit Transitional Member Cache Valley Member and Tscl Lower conglomerate unit Skyline Member 8 Field trip stop Wasatch Formation

Ttc-cv Tcv Ts

Clifton detachment fault

Q

Quaternary-Tertiary gravel

Ttc

Zp

K

St

Q

Clarkston strand of the West Cache fault

QTg

Tw

Tscu C-Zb

O-D

36

E 15

T RSTA

INTE

Henderson Creek quadrangle

Tcv

Area of Figure 7

Y

Tscu C-O

C-O

Ttc

EE

C-OTscl

Ttc-cv Q R

Tscl

O-D

C

Tcv

Deep Creek fault zone

ent segm ntain tch fault Mou sa e Wa of th

Tcv

C-Zb

HW

Tcv

C-Zb

9

Ttc

C-O

N

O-D

C-O

TO

O-D

Tcv

Ttc

O-D

Tcv

Tcv

Q

O-D

Miocene-Pliocene Salt Lake Formation

O-D

Tcv

Tcv C-Zb

8 Ttc

O-D

1 Q

Ttc-cv

Q

Tcv

C-O

Q

WOODRUFF/SAMARIA EXIT

Oxford Ridge anticline

Zp Ttc

Tcv

Tcv

TcvTscl

Ttc

Ttc

C-O

Q

Ttc

ES W

ston

Clark Tcv

O-D

C-O

ON NY CA

C-O

5

Zp Q

C-O

ON ST

t en lt gm u se fa ty tch Ci sa ad Wa al M the of

O-D

Tw

Tcv

7

Q Q

O-D C-O

Tw O-D

Tnc

Q

Ttc

Tcv

Ts

Tcv

Ttc

Ttc

Tw Ts

Zp

Ttc

Tcv

ll fault

Ts

Tnc

C-O

Tcv

Y3

6 ER VO IR

Red Kno

e ylin n Sk grabe lf ha NGE Tw

Q

Ttc C-O

QTg

HW

RE S

Tw

Q

y

C-Zb

Tcv

D RA O-D

Tnc

C-O

WE

MALA

O-D

Q

Q Q

W ES TO N

O-D

Ts

C fa em ul et t er

C-O

Q

Q

Tw

on

C-O

Tcv

QTg

C-O

Ts

ift

C-Zb

C-O

QTg

C-O

O-D

O-D

Q

C-Zb

Cl

Zp

Q

O-D

O-D

C-O

MALAD VALLEY

Ts

Ttc

Ttc

C-O

CLIFTON

Tnc

Clifton Basin

QTg

QTg

Ttc

Ts

C-O

Ttc

QTg QTg

Tw C-O

Tcv

Q

Q

C-Zb

Ts

Ts

Q

O-D

QTg Q

Red

15 Malad City segment of the Wasatch fault

C-Zb

Ts

Tw

C-Zb

SKYLINE ROAD

TE STA

Tw

C-Zb

Ttc QTg

su ep C b- re ba e sin k TW OM

Ts

C-O

Ttc

Tcv

De

Tcv

ne ek fault zo Deep Cre

C-O

Zp

C-Zb

O-D

Q

QTg

RST

O-D

Zp

Ttc

CLIFTON HO

Tw

O-D

Ttc

Ttc

lt 3 fau ford Ox onayt

Tcv

4

C-Zb

tD

MALAD CITY

Q

Tcv

Oxford Ridge anticline

Eas

Pocket Basin

Ttc O-D

Ttc Zp

Y2 HW

DEEP CREEK RESERVOIR

Tcv

Tcv

O-D

Zp

NGE

Ts

Q Zp

West Dayton-Oxford Fault

Ttc

Q

K RA

Q

Q

Q Zp

NOC

O-D

Tcv

Ts

Tw Ts

Q

Zp

Ttc

Q

Tcv

BAN

Ttc

Q

Q

Tertiary intrusion Zp

E RIDG

Zp

QTg

HWY 36

Q

Q

Weston Canyon quadrangle

Q

Zp

Tcv

Neoproterozoic Pocatello Formation high-angle normal fault major Basin-and-Range fault low-angle normal fault plunging anticline plunging syncline monocline zeolitized tuffaceous marker bed

O-D

Ordovician, Silurian, and Devonian strata

C-O

Cambrian and Ordovician strata

C-Zb

Cambrian and Neoproterozoic Brigham Group

conglomerate beds

approximate attitude of bedding outline of persistent structural trend

42˚ 00’ N

Q

O-D

RD OXFO

Tnc

Tcv

East DaytonOxford fault 42˚ 07’ 30” N

Ttc

Q

112˚ 00’ W

Zp

Clifton quadrangle

Ttc

42˚ 15’ N

A.N. 112˚ Steely et al. 07’ 30” W

Malad City East quadrangle

Dayton-Oxf ord Fault

174 112˚ 15’ W

Figure 4. Simplified geologic map of Clifton, Henderson Creek, Malad City East, and Weston Canyon 7.5′ quadrangles. See Fig. 2 for location. Compiled from Evans et al. (2000), Carney (2002), Long (2004), and Steely and Janecke (2005).

Evolution of a late Cenozoic supradetachment basin A) Low-angle normal fault models

175

B) Excision by master fault Breakaway

Inactive, older normal fault

Original low-angle

Active normal fault

1

Future normal fault

2 Breakaway Future position of fault

Domino model

3 Isostatic folding

4 Breakaway

Hinge

Rolling hinge model

Hinge

Excised fault block

Secondary breakaway

Present position of fault

5

portion

total slip cision ue to ex p of sli d

Excised fault block

Figure 5. (A) Schematic diagrams showing three models of extensional low-angle normal faults. (B) Schematic diagram showing excision: 1—original geometry; 2—slip on detachment fault; 3—doming; 4—creation of a new low-angle listric breakaway. Higher angle faults are cut by the low-angle fault; 5—continued slip juxtaposes synrift deposits in the hanging wall with rocks in the footwall of the detachment fault for the first time. Excised fault block is shaded. Modified from Carney and Janecke (2005).

was important during the evolution of the Bannock detachment fault (Carney and Janecke, 2005). The original three-dimensional geometry of low-angle normal faults is critical to understanding how folds and corrugations evolve within the fault system and how hanging wall deformation is affected by fault geometry. Because some detachment faults slip within a few kilometers of Earth’s surface (Smith, 1984; Otton, 1995; Axen et al., 1999; Carney and Janecke, 2005), ramp-flat and lateral ramp geometries can significantly affect the hanging wall thickness and consequently may play a pivotal role in determining the deformational style of the hanging wall. These fault geometries may also affect sedimentation patterns in the overlying supradetachment basin. An important feature of low-angle normal fault systems is the formation of supradetachment basins and their associated fill. These basin-fill deposits allow us to reconstruct the timing of deformation events through careful analysis of stratigraphy,

sedimentology, and structural geology, coupled with geochronology. Recent work (VanDenburg et al., 1998; Janecke et al. 2003; Janecke and Blankenau, 2003; Janecke, 2004), building on earlier work by Friedmann and Burbank (1995), has shown that distinct stratal architectures result from the two-stage evolution of many supradetachment basins. The first phase, called the translation phase, is marked by conformity to subtle-fanning dips within widespread basin fill deposited during simple translation. The second phase of evolution begins as the basin breaks up internally into smaller half graben above the master detachment fault. This phase is commonly marked by angular and progressive unconformities, dramatic decreases in basin size, the potential for more lateral variations in depositional systems relative to the translation phase, and the presence of clasts recycled from translation-phase deposits. The recycled clasts are particularly significant because they demonstrate that some areas of the previously intact basin were uplifted and eroded during breakup

176

A.N. Steely et al.

(Janecke and Evans, 1999; VanDenburg et al., 1998; Janecke et al. 2003; Janecke and Blankenau, 2003). STRATIGRAPHY OF STUDY AREA Bedrock strata in the study area consist of the Neoproterozoic Pocatello Formation, Neoproterozoic to Cambrian Brigham Group rocks, and Cambrian to Devonian miogeoclinal carbonates and shales deposited within the Cordilleran miogeocline (Fig. 6). Silurian to Permian rocks crop out to the west of the study area in the Samaria Mountains and to the east in the Bear River Range (Platt, 1977). Any late Paleozoic to Mesozoic units that once covered the area were eroded from the Cache-Pocatello culmination before Eocene time (Fig. 2). The Paleocene to Eocene Wasatch Formation, late Miocene to Pliocene Salt Lake Formation, and Pliocene-Pleistocene(?) piedmont gravel deposits overlie the Paleozoic and older rocks along an angular unconformity (Fig. 6). The distinctive lithologies of the pre-Tertiary bedrock in the study area allowed reconstructions of specific source areas for the late Cenozoic basin-fill deposits. Quaternary lacustrine and near-shore deposits of the Lake Bonneville cycle cover most areas below ~1,560 m (5,120 ft) in elevation. Pre-Cenozoic Stratigraphy Neoproterozoic Pocatello Formation The Neoproterozoic Pocatello Formation is divided into three distinct members: the Scout Mountain Member, the Bannock Volcanic Member, and an unnamed upper member (Fig. 6) (Link, 1982a, 1982b; Link et al., 1993; Smith et al., 1994). Together the Pocatello Formation has an aggregate thickness of 450 m just north of the study area (Link, 1982a). All of the members have been metamorphosed to greenschist facies (Link, 1982a) and are foliated in the Clifton and Weston Canyon quadrangles. Only the Scout Mountain Member is present within Weston Canyon quadrangle, where it crops out from north of Fivemile Canyon to the southern end of Rattlesnake Ridge in the south and attains a minimum thickness of ~350 m (Fig. 4). The Bannock Volcanic Member is exposed along Oxford Ridge in the Clifton quadrangle (Carney, 2002). 709 Ma zircons in a metatuff in the Clifton quadrangle indicate that the Scout Mountain Member is Sturtian in age (Fanning and Link, 2004). Within the Weston Canyon quadrangle, the Scout Mountain Member consists of green to green-brown siltstone, fine- to coarse-grained sandstone, and diamictite, with local interbeds of 1–3-m-thick tan to tan-gray limestone marble along the uppermost exposures of Rattlesnake Ridge. The Scout Mountain Member is composed dominantly of siltstone and rare to locally abundant matrix-supported, pebble to cobble conglomerate. Clasts dispersed in the diamictite include metavolcanics, quartzites, and granitoids. In the footwall of the Bannock detachment system, weaker pebbles (metavolcanics) are stretched where a pervasive weak to strong foliation is also present, and sometimes preserve two lineations. Some areas have two foliations at acute angles to one another.

Because foliation has overprinted primary bedding within most of the footwall rocks, understanding the relationship between these two fabrics is crucial in constraining the cut-off angle between footwall rocks and the Bannock detachment fault. Several locations within the Weston Canyon quadrangle document primary bedding of the Pocatello Formation and confirm that it is parallel to subparallel with foliation fabrics where both are observed on the west limb of the Oxford Ridge anticline. Outcrops at A on Figure 7 expose well-bedded siltstone and sandstone, which are overprinted by bedding-parallel foliations. The north end of Rattlesnake Ridge (B on Fig. 7) exposes thin- to medium-bedded siltstone, sandstone, and small pebblegranule conglomerate with bed-parallel foliations. Along the southern part of Rattlesnake Ridge, the outcrop pattern of a thin marble bed is parallel to foliation in the underlying and overlying diamictite (C on Fig. 7). Bedding parallel foliations like these are documented at the base of the Paris-Willard thrust sheet along structural strike to the south (Yonkee et al., 1997). Late Neoproterozoic-Cambrian Brigham Group and Paleozoic Strata The late Neoproterozoic to Cambrian Brigham Group contains ~1.6 km of quartzite, pebbly quartzite, sandstone, vitreous quartzite, and micaceous argillite (Fig. 6) (Link et al., 1987; Carney et al., 2002). Only the upper five of eight units of the Brigham Group crop out within the Weston Canyon quadrangle (Fig. 6; Mutual Formation, Rocky Peak Phyllite Member, Camelback Mountain Quartzite, Windy Pass Argillite, and Sedgwick Peak Quartzite). Cambrian to Devonian rocks of the Cordilleran miogeocline crop out throughout the four quadrangles and have a composite thickness of ~2.9 km (Fig. 6) (Biek et al., 2003; Carney et al., 2002; Long, 2004). These rocks are dominantly composed of limestone, dolostone, calcareous to non-calcareous shale, and lesser sandstone and quartzite. The distinctive lithologies of four main groups of pre-Tertiary rocks (Pocatello Formation, Brigham Group, middle to lower Paleozoic rocks, and upper Paleozoic rocks) allow us to determine the source of conglomerates and to reconstruct the paleogeography of the Tertiary deposits. Within the Paleozoic rocks, several key lithologies allow us to further refine the approximate stratigraphic level exposed. Some of these key lithologies are a vitreous bright white quartz arenite of the Swan Peak Formation, which is readily distinguished from the red to purple quartzites of the Brigham Group quartzite. Black chert from the Garden City Formation, tan chert from the St. Charles Formation, and reddish-orange–weathering sandy limestone of the Oquirrh Formation are also distinctive. Cenozoic Stratigraphy Wasatch Formation The Wasatch Formation is an early to middle Eocene deposit of the Sevier fold-and-thrust belt in the Idaho-Wyoming-Utah region (Coogan, 1992; Oaks and Runnells, 1992;

Evolution of a late Cenozoic supradetachment basin

Missing Paleozoic stratigraphy

?

Tnc

New Canyon 6000 ft Member

2000 m

1500 m 4000 ft 1000 m 2000 ft

Late Miocene

Ttc

Third Creek Member 0

Ttc-cv

?

?

Eocene

Devonian

Tcv

Cache Valley Member

Ts

Skyline Member

Tw

Wasatch Formation

Dj Dwc

Jefferson Formation Water Canyon Formation Undivided upper OrdovicianSilurian including the Fish Haven Dolomite Swan Peak Quartzite

Ogc

Garden City Formation

Cambrian

Cbo

Bloomington Formation

Cbl

Blacksmith Formation Ute Formation

Salt Lake Formation

Neoproterozoic Pocatello Formation

Third Creek Member (Ttc) and Tscu of Long (2004) New Canyon Member (Tnc)

Neoproterozoic-Cambrian Brigham Group

? Conglomerates of the Cache Valley Member (Tcv) and Tscl of Long (2004)

?

Upper Paleozoic rocks exposed west of the Malad Range

Osp

Cu

Skyline Member (Ts)

0

OCsc St. Charles Formation Cwc Worm Creek Formation Cn Nounan Formation

?

?

Ordovician

SOu

500 m

Third Creek-Cache Valley Transitional Member

?

Provenance of Tertiary conglomeratic basin fill deposits

Wasatch Formation (Tw)

Salt Lake Formation The Salt Lake Formation is the sedimentary product of extension in the northeast Basin and Range province (Janecke and Evans, 1999; Oaks et al., 1999; Goessel et al., 1999; Janecke et al., 2003; Kruger et al., 2003). A strong spatial association exists between the Salt Lake Formation and highly extended areas above detachment faults (Carney and Janecke, 2005). Within the Bannock and Malad ranges, all but the basal Skyline Member of the Salt Lake Formation has been interpreted as synextensional basin fill deposits above the Bannock detachment fault system (Janecke and Evans, 1999; Janecke et al., 2003). The Skyline Member filled early, pre-detachment half-graben (Janecke et al., 2003; Long, 2004). It was overlain by the Cache Valley Member of the Salt Lake Formation, which filled one continuous supradetachment basin above the Bannock detachment fault. The Valley fault in the Portneuf Range, its along-strike continuations, and the ancestral–East Cache fault are interpreted as the breakaway for the Bannock detachment system (Fig. 1) (Janecke and Evans, 1999; Janecke et al., 2003; Eversaul, 2004). After a period of translation this large basin broke up internally. Smaller half graben formed above the detachment fault, and older synrift strata were uplifted, eroded, and redeposited as clasts in younger synrift strata of the Third Creek and New Canyon members of the Salt Lake Formation (Janecke et al., 2003). The Salt Lake Formation consists of four members (Figs. 6 and 8), and three of these crop out in the Weston Canyon and Henderson Creek quadrangles (Skyline, Cache Valley, and Third Creek members). The maximum intact and unfaulted exposure of Salt Lake Formation is ~2.5–2.7 km thick west of Rattlesnake Ridge in the Weston Canyon quadrangle (Fig. 7). The Salt Lake Formation has a composite thickness of >4.3 km in the Malad and SE Bannock ranges. Skyline Member. The Skyline Member is the basal unit within the Salt Lake Formation, and lies along angular uncon-

Pliocene

Definite provenance

QTg QTrg Quaternary-Tertiary gravel Quaternary-Tertiary roundstone gravel

? ?

Possible provenance

?

Oaks et al., 1999; Long, 2004). The unit is considered syntectonic with late main-stage movement on the frontal Hogsback thrust in western Wyoming (DeCelles, 1994) but also laps across the entire thrust belt as far west as the Malad Range. It was deposited in both N-S- and E-W–trending Eocene grabens (Oaks and Runnells, 1992; Long, 2004) above the ParisWillard thrust sheet after movement ceased on this thrust (Coogan, 1992). The Wasatch Formation either unconformably overlies or is faulted against Cambrian through Devonian rocks in the Malad Range and Oxford Ridge area (Fig. 6) (Carney, 2002; Long, 2004). It generally consists of red, moderately to poorly consolidated, matrix-rich, pebble to boulder conglomerate with a thickness of <189 m in the Henderson Creek quadrangle. It is interpreted as a braided stream and locally-derived alluvialfan deposit that shows evidence for syndepositional tilting and faulting in an asymmetric south-tilted half graben (Fig. 8) (Long, 2004). The Wasatch Formation is dominated by locallyderived middle Paleozoic-age clasts (Fig. 9).

177

Cl Cwp CZcm

Langston Formation Windy Pass Argillite and Sedgewick Peak Quartzite Camelback Mountain Quartzite

Zrp

Rocky Peak Member of Camelback Mountain Quartzite Zi Mutual Formation Inkom Formation Zcc Caddy Canyon Quartzite Zpc Papoose Creek Formation Zm

Zpu

Pocatello Formation upper member

Zps

Pocatello Formation Scout Mountain Member

Zpb

Pocatello Formation Bannock Volcanic Member

Zps

Pocatello Formation Scout Mountain Member

Figure 6. Generalized stratigraphic column of strata exposed in the field trip area and plot showing the provenance of Tertiary conglomeratic basin fill deposits. Note the overall unroofing sequence, absence of Pocatello Formation in basin fill, and recycling of Tertiary rocks in the Third Creek Member.

A.N. Steely et al.

178

Cbl Cbl Cbl

CZcm Cbl Cbl

CZcm Cbl

CZcm

Cbl

CZcm

CZcm CZcm

CZcm

CZcm CZcm CZcm CZcm

Cbl

Cbl

Cbl

CZ cm

Figure 7. Simplified geogic map of the northeastern Weston Canyon and southeastern Clifton quadrangles. Note the concordance of the gravity saddle with the trend of the lateral ramp in the detachment fault. BDF—Bannock detachment fault system (includes the Clifton and Bannock faults); ORA—Oxford Ridge anticline. Modified from Steely and Janecke (2005) and Carney (2002).

CZcm

sandstone conglomerates of Third Creek Member

Transitional Member

Cache Valley Member

Skyline Member

Wasatch Formation

Ttc-cv

Tcv

Ts

Tw

Miocene-Pliocene Salt Lake Formation

age or age range based on tephra correlation

39

?

ol l fa

ult

Ts

Tw

10.27 Ma ( Ar/ Ar)

Tcv

40

Ttc

Tnc

~10.13 Ma

?

Kn

lower Paleozoic rocks

Tw

~8.3 Ma

5.1 Ma ~7.9 Ma

QTg

Deep Creek half graben

NW

10.13-8.5 Ma

conglomerate

?

?

?

?

?

angular unconformity or disconformity

2m 3 ay m.y be . o m fs iss ect in ion g

tephra ( sample location, * is a correlation in conflict with detrital zircons)

tephra correlation tie line

gravel

Third Creek Member

Ttc

limestone

New Canyon Member

?

~9.6 Ma

faulted by Bannock detachment fault

Tcv

Ttc-cv

Ttc

Ttc

Faulted

?

East

?

?

500 meters

Growth folds

Basal conglomerate beds restricted to growth synclines

Weston Canyon quadrangle

invertebrate fossil

Tnc

approx. 500 m

? ?

Quaternary-Tertiary gravel

R ed

C-D rocks

Tw

11.93 Ma*

Ts

C-D rocks

Tw

O-D rocks Willow Spring fault

Red Knoll and Trespass faults

C-D rocks

Tw

9.81 Ma

10.02 Ma*

9.8 Ma

9.38 Ma

9.59 Ma 9.67 Ma

Tscu

C-O rocks

Willow Spring fault

O-D rocks

Tscl

10.20 Ma

Tcv-tufa bearing facies

Faulted or eroded 9.3 Ma

Tcv-micrite bearing facies

9.16 Ma

Henderson Creek quadrangle

C-O rocks

Steel Canyon and Burnett Canyon faults

Z- C rocks

Z-T rocks

SW

Figure 8. Simplified Tertiary stratigraphy of the Salt Lake and Wasatch Formations in the Bannock and Malad ranges, SE Idaho. Fence diagram looks east toward the Weston Canyon quadrangle. Tie lines are based on lithologic similarities, tephra correlations, and one 40Ar/39Ar date. Note the significant lateral changes in the thickness of the Skyline, Third Creek, and New Canyon members. The Cache Valley Member is much more uniform in thickness and occurs at all locations. The Weston Canyon section may contain the hypothesized missing stratigraphy in the Malad City East quadrangle. Deep Creek half graben stratigraphy modified from Janecke and Evans (1999). Henderson Creek stratigraphy modified from Long (2004). Tephra correlations are from collaborations with Michael Perkins and Barbara Nash.

500 meters

QTg

Evolution of a late Cenozoic supradetachment basin 179

A.N. Steely et al.

180 100

100

New Canyon Member n=7 sites (Tnc) <4.4 Ma

60 87%

Skyline Member (Ts) ~12 to ~10 Ma

80

n=26 sites

69% Percent

Percent

80

40

60 40

20

19%

20

10%

9% 3%

0 Lower-Middle Paleozoic carbonate, chert, and quartzite

Upper Paleozoic sandstone, siltstone, and limestone

Recycled Tertiary tuffaceous siltstone and limestone

52%

Third Creek Member (Ttc) and Upper Conglomerate (Tscu) of Long (2004) >9.6 to <4.4 Ma n=68 sites

40

80 Percent

80 Percent

CambrianNeoproterozoicBrigham Group quartzite

100

100

60

<3%

0 CambrianNeoproterozoicBrigham Group quartzite

Lower-Middle Paleozoic carbonate, chert, and quartzite

16%

Wasatch Formation (Tw) Eocene n=9 sites 86%

60

20

14%

0

0 CambrianNeoproterozoicBrigham Group quartzite

Lower-Middle Paleozoic carbonate, chert, and quartzite

Upper Paleozoic sandstone, siltstone, and limestone

Recycled Tertiary tuffaceous siltstone and limestone

CambrianNeoproterozoicBrigham Group quartzite

100

Percent

Recycled Tertiary tuffaceous siltstone and limestone

40

32%

20

80

Upper Paleozoic sandstone, siltstone, and limestone

94%

n=21 sites

60 40

Conglomerates of the Cache Valley Member (Tcv) and Lower Conglomerate (Tscl) of Long (2004) 10.2 to 9.2 (?) Ma

Lower-Middle Paleozoic carbonate, chert, and quartzite

Upper Paleozoic sandstone, siltstone, and limestone

Recycled Tertiary tuffaceous siltstone and limestone

Figure 9. Combined clast count data from conglomeratic basin-fill deposits of the Clifton, Henderson Creek, Malad City East, and Weston Canyon 7.5′ quadrangles (see Fig. 4 for quadrangle locations). Note the unroofing sequence shown by upsection increases in Brigham Group quartzites with a simultaneous decrease in Paleozoic clasts. Also note the sudden influx of recycled clasts of tuffaceous and calcareous Salt Lake Formation during deposition of the Third Creek Member. Clast count data is from this study (n = 21 for unit Ttc; n = 6 for Tcv conglomerates); Janecke and Evans (1999), Carney (2002), and Long (2004) (n = 58); n—number of clast counts in each chart (each count represents either 50 or 100 clasts at a single location).

20 4% 2%

0 CambrianNeoproterozoicBrigham Group quartzite

Lower-Middle Paleozoic carbonate, chert, and quartzite

Upper Paleozoic sandstone, siltstone, and limestone

Recycled Tertiary tuffaceous siltstone and limestone

formity on the Wasatch Formation and lower Paleozoic rocks. It is preserved in the western Deep Creek half graben within the Skyline subbasin at the southern end of the Malad City East quadrangle and the north end of the Henderson Creek quadrangle (Fig. 4) (Janecke and Evans, 1999; Long, 2004). The Skyline Member is up to 685 m thick in the Henderson Creek quadrangle (Fig. 8), and most beds are >30 cm thick. The unit consists of poorly sorted, tuffaceous, pebble to cobble conglomerate with interbedded tephra beds and locally interbedded lacustrine limestone (Janecke and Evans, 1999; Janecke et al., 2003; Long, 2004). Tephra beds are light-colored, up to 5 m thick, and commonly contain angular gray and black glass shards. Clasts within conglomeratic intervals are composed dominantly of lower and middle Paleozoic carbonate, chert,

and minor quartzite but include some upper Paleozoic units as well. However, light-colored tuffaceous to calcareous siltstone clasts are also present within the Skyline Member. These clasts are either rip-ups from Miocene limestone and tephra units within the Skyline Member of the Salt Lake Formation or were derived from surrounding areas where Salt Lake Formation deposits were being uplifted and recycled. A 90-m-thick section of crystalline and micritic white limestone is present near the base of the member in the Henderson Creek quadrangle. This limestone deposit suggests that earliest deposition of the Salt Lake Formation may have occurred in small lacustrine basins. Conglomerates of the Skyline Member are interpreted as alluvial-fan deposits that filled a local east-tilted half graben (Janecke et al., 2003; Long, 2004).

Evolution of a late Cenozoic supradetachment basin Tephra correlations from the Skyline Member reported in Long (2004) indicate an age of 11.93 ± 0.03 Ma near the base and 10–11 Ma near the top, but these ages are being reevaluated in light of a large population of 14–15 Ma detrital zircons in a slightly reworked tephra near the top of the Skyline Member. The Skyline Member may be correlative with the Collinston Conglomerate farther south (Goessel et al., 1999). Both deposits are coarse basal units of the Salt Lake Formation and fill local half graben above west-dipping normal faults. Cache Valley Member. The Cache Valley Member overlies the Skyline Member along a gradational to sharp contact in the Skyline half graben (Janecke and Evans, 1999; Long, 2004). Elsewhere it overlies lower Paleozoic rocks and the Wasatch Formation along a slight angular unconformity (Fig. 8). It is exposed primarily on the west and south sides of the Deep Creek half graben and on the east and west sides of Oxford and Rattlesnake Ridges in the hanging wall of the Bannock detachment fault (Fig. 4). It is at least ~617 m thick based on map measurements west of Rattlesnake Ridge. This thickness is similar to a thickness of 610 m of lacustrine beds in the Henderson Creek quadrangle (Long, 2004), and a minimum thickness estimate of 600 m in the Clifton quadrangle (Carney, 2002). Oaks (2000) estimated >746 m thickness of the Cache Valley Member along the northern Wellsville Mountains, and Goessel et al. (1999) and Biek et al. (2003) showed >975 m between the base and their overlying oolitic subunit in the Junction Hills and southern Clarkston Mountain. The relatively uniform thickness of the Cache Valley Member in the Malad and SE Bannock ranges contrasts sharply with the abrupt lateral thickness changes of the overlying and underlying members of the Salt Lake Formation (Fig. 8). A rhyolite tuff at the base of the Cache Valley Member in the Malad City East quadrangle is a distal exposure of the tuff of Arbon Valley dated at 10.21 ± 0.03 Ma (Morgan and McIntosh, 2005). This tuff has a distinctive mineralogy of smoky quartz, sanidine, plagioclase, and biotite crystals and provides a useful marker when present. Long (2004) observed the tuff of Arbon Valley in the basal Cache Valley Member and obtained a tephra correlation of 9.16 ± 0.03 Ma from high in the lacustrine Cache Valley Member in the Henderson Creek quadrangle. This is the youngest preserved Cache Valley Member in our area (Fig. 8). Elsewhere, Cache Valley lithofacies had been replaced by conglomerate-bearing Third Creek lithofacies by this time. Within the Weston Canyon quadrangle the Cache Valley Member consists of interbedded tan-white limestone to silty limestone with white to green tuffaceous mudstone, siltstone, and sandstone. Lesser calcareous mudstone and siltstone, silicified laminated limestone, rare fine- to medium-grained sandstone and discrete beds of poorly sorted granule to rare small boulder conglomerate crop out in the quadrangle. Bedding within the unit ranges from laminated to thick and poorly bedded. A slight overall upsection increase in sandstone content within the Cache Valley Member occurs west of Rattlesnake Ridge. Rare conglomerate beds in the Cache Valley Member are localized within the Weston Canyon and Henderson Creek quadrangles near the base

181

of the unit. Within the Weston Canyon quadrangle, one to three conglomerate beds extend from the north edge of the quadrangle to the southern end of Rattlesnake Ridge across a distance of ~7 km (Fig. 7). These beds are 1–4 m thick, composed of poorly sorted, matrix- to clast-supported, subangular to well-rounded granules to cobbles and rare small boulders of lower-middle Paleozoic carbonate, lesser Paleozoic quartzite, locally abundant black and tan chert, and green quartzite. This green quartzite is probably stained Eureka Quartzite Member of the Ordovician Swan Peak Formation because the other possible source of green quartzite, in the upper Brigham Group, is interbedded with distinctive purplish to pink quartzite that is not present in the conglomerate beds of the Cache Valley Member (Figs. 9 and 10A). The conglomerate matrix is dominantly light-colored micrite mud, although some beds have a calcareous sand groundmass. These beds appear to overlie tuffaceous siltstone and sandstone along gradational to sharp contacts and are overlain along sharp contacts with micritic limestone or silicified laminated limestone. Just east of Fivemile quarry in the Weston Canyon quadrangle, conglomerate beds constitute ~5% of the lower ~220 m of the Cache Valley Member and pinch out north and south along strike into white to tan tuffaceous to calcareous siltstone, sandstone, or mudstone. Conglomerate beds persist 100–500 m along strike. Lateral facies changes in the Cache Valley Member in the Henderson Creek quadrangle reflect a southwestward shoaling toward an intrabasinal fault block. Exposures in the western part of the Henderson Creek quadrangle are dominated by ledgeforming tufas interbedded with micritic limestone, reworked tuffaceous siltstone and sandstone, and primary tephra beds. The tufa-facies interfingers with and passes laterally eastward into micrite-dominated limestone and tuffaceous rocks. Micritic limestone is white to light gray, thin- to medium-bedded, and also contains thin silicified stringers. Both the micrite- and tufa-bearing facies locally contain ostracodes, gastropods, and pelecypods. These lithologies interfinger westward with a ~5.7-kmlong lens of syntectonic conglomerate shed from the then active Steel Canyon fault in the upper 60% of the Cache Valley Member (Long, 2004). The oldest conglomerate is 9.67 ± 0.09 Ma (Long, 2004), approximately the age of the oldest conglomeratic Third Creek Member farther to the north (Janecke et al., 2003). These data show that Cache Valley lithofacies overlap in age with Third Creek lithofacies in the Henderson Creek quadrangle (Fig. 8). In the Weston Canyon, Clifton, and Malad City East quadrangles, tuffaceous rocks of the Cache Valley Member have been altered to clays and zeolites to varying degrees (Janecke et al., 2003). These rocks are most commonly greenish, but may also locally be off-white, yellow, and tan, and are locally much more indurated than unaltered rock. A hackly fracture pattern is common. These highly altered rocks are in marked contrast to silvery, light gray primary tephras of the Third Creek Member (discussed below), which are much less altered and are poorly consolidated (Janecke et al., 2003). Interpretation. We interpret the Cache Valley Member to represent widespread lacustrine deposition during “translation-

A.N. Steely et al.

182 A

B

CZb

Tsl

Tsl

CZb

CZb Tsl

Tsl Tsl

CZb Tsl q

Tsl

Tsl

Tsl

q

Tsl

Tsl

Tsl Tsl

Tsl

Tsl q

Tsl

Tsl

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Tsl

Tsl

Tsl

Tsl Tsl

q

CZb

Tsl Tsl

Tsl

Tsl

Tsl

Tsl

CZb CZb

CZb

Tsl

C Figure 10. (A) Conglomerate bed in lower Cache Vallye Member at Stop 5. Note angularity and micrite supported clasts of dominantly gray Paleozoic rocks (inset). Green clasts are stained quartzite of the Ordovician Swan Peak Formation (q in photo). (B) Conglomerate of the Third Creek Member. Recycled clasts of the Salt Lake Formation (green; Tsl) outnumber quartzite clasts from the Brigham Group in this bed (red; CZb). Clasts are much more rounded than in conglomerates of the Cache Valley Member. (C) Exposure of the Bannock detachment with down-dip slickenlines developed on brecciated Cambrian-Neoproterozoic Camelback Mountain Formation. Fault places lower Cache Valley Member against Cambrian and Neoproterozoic rocks. Slickenlines trend WSW (Fig. 12D). View is to the east.

phase” westward motion of the relatively coherent hanging wall of the Bannock detachment fault (Fig. 11). This member is present in every part of the Salt Lake basin and is relatively uniform in its maximum thickness across lateral distances of ~20–30 km. Conglomerate beds in the lower Cache Valley Member in the Weston Canyon quadrangle are interpreted as distal alluvial fan deposits that interfinger with near shore lacustrine deposits. The micrite matrix within the conglomerates suggests that they were deposited during a time of reduced ash input to the lacustrine system. The presence of pebble conglomerates in the lake suggests that there were highlands nearby, likely to the east, that were eroding during deposition of the lower Cache Valley Member deposition. These highlands may have been produced either by slip on an early intrabasinal breakup fault or there may have been remnant topography from older faults. The disappearance of these conglomerates upsection indicates that: (1) the remnant topography was fully eroded, (2) the intrabasinal fault ceased movement,

or (3) progradation of conglomerates into the “Cache Valley lake” is controlled by the amount of ash covering local topography. Relationships in the Henderson Creek quadrangle suggest that intrabasinal fault-bounded horst blocks began to produce subbasins within the large Cache Valley lacustrine system as early as 9.6 Ma. Despite localized conglomeratic input from the footwall of the Steel Canyon fault, the lake filled mostly with ash from the Yellowstone hotspot and with micrite and tufa. Carbonate deposition was localized in the shallower parts of the lake between major volcanic eruptions, with tufa near the shoreline and micrite in more offshore positions (Long, 2004). Facies patterns show that the apparent change in water chemistry was accompanied by shoaling to the south and west. The lake beds locally passed laterally into syntectonic intrabasinal alluvial-fan conglomerates in the Henderson Creek and eastern Weston Canyon quadrangles. Persistent lacustrine conditions may reflect a relatively stable sill and steady input

Evolution of a late Cenozoic supradetachment basin

A

B

Deposition of the Cache Valley Member of the Salt Lake Formation ~10.2 Ma to 9.6 Ma 50% extension restored

183

Deposition of the Third Creek member of the Salt Lake Formation (~9.6 Ma to <4.4 Ma)

Fu

on

or eP tur

tt Co

mixed facies

ng e

rr Na

rgi

s ow

ma

Malad Valley subbasin

Ra

Saline/alkaline lake

ida

rn ste we

Swan Lake subbasin

Deep Creek/Oneida saline/alkaline lake

ne

Eroding fault block

uf

/O ey all dV

st -ea

Alluvial fan and fan/ delta deposits

tne

o wo

s? nd lt hla au hig ef ng ch Ca

lley Va

i erg em

25% extension restored

ain er t

nc

Future Bear River Range

nu

Deep Creek subbasin

Freshwater/brackish lake/marsh Future normal fault

Utah

SW Freshwater lake SM margin uncertain

MV

saline/alkaline lake DC

CV/ON

NE

Active normal fault

Idaho Utah

Stream

ult

East Cache fa

Idaho

Future Portneuf Range

deeper lake

SW

NE

MV

Clifton faul

DC

t

SL

CV/ON

Extensional horses Future Clifton horst

Figure 11. Reconstructed paleogeographic maps and cross sections of the northern Cache Valley area between the Samaria Mountains and the Bear River Range. Faults that developed during the next episode of deformation are shaded and dashed. (A) Initiation of slip on the detachment system and development of freshwater–saline-alkaline lake conditions with small intrabasinal highlands (Cache Valley Member). (B) Hanging wall breakup of the detachment system along southwest-dipping normal faults. Smaller half grabens formed; older Tertiary basin fill of the Salt Lake Formation was uplifted and recycled into younger freshwater lacustrine, nearshore, and fluvial deposits (Third Creek Member). Modified from Janecke et al. (2003).

of water into the lacustrine system. Overall, the “Cache Valley lake” system was widespread and was the only laterally-continuous depositional system in the Salt Lake Formation. Third Creek–Cache Valley Transitional Member. The transitional member was only mapped separately in the Weston Canyon quadrangle where the presence of two marker tuffs helps to distinguish it (Figs. 4 and 7). There it lies conformably above the Cache Valley Member. Overall, clastic interbeds within this unit coarsen upward from the tuffaceous and calcareous Cache Valley Member into the conglomerate-bearing Third Creek Member (Fig. 8). The base of the transitional unit is placed at the first appearance of red to pink quartzite grains (Brigham Group) in sandstone. This change coincides with the lower of two prominent tuffaceous siltstone marker beds, except locally where lithofacies of the Cache Valley Member and the transitional member interfinger (D on Fig. 7). The top of the transitional unit is defined by the lowest persistent pebble to cobble conglomerate of the Third Creek Member. These lens-shaped conglomerate beds crop out just below the upper prominent tuffaceous siltstone marker bed (Figs. 4 and 7). Where conglomerates are not present, the contact is placed at the base of the upper marker bed.

The transitional unit consists of tan-gray limestone, gray tufa, fine- to coarse-grained sandstone, white to light-green tuffaceous siltstone and sandstone, and poorly exposed white- to light-green tephra. Red quartzite and tuffaceous sand grains are derived from the Brigham Group and Salt Lake Formation, respectively. Rare flat pebble conglomeratic sandstone with Cache Valley Member clasts crops out locally. This unit marks the first appearance of widespread gray tufa that locally preserves tufa heads with a spongy texture. The limestones in this unit are petroliferous when fresh or exposed to HCl, are usually laminated to thinly bedded, locally oolitic, and locally contain whole gastropod shells and shell fragments in oolite-rich beds. The thickness of this unit varies along strike west of Rattlesnake Ridge where several large west-plunging folds are present (Fig. 7). Based on map measurements, the transitional member thins from 368 m in the core of the southern major syncline to 198–217 m on the two flanking anticlines. Just south of Fivemile Creek, a smaller anticline-syncline pair complicates the hinge zone of a second enveloping syncline (Fig. 7). The lowest conglomerates of the Third Creek Member are localized in these enveloping synclines and pinch out over the crests of adjacent anticlines

184

A.N. Steely et al.

(Fig. 7). Outcrop patterns and map relationships show that these folds deform the lower transitional member and the underlying Cache Valley Member equally. Strata of the lower Third Creek Member are less deformed than the transitional member. Interpretation. The transitional member marks a change from widespread, open water lacustrine deposition of the Cache Valley Member into shallower water, lacustrine, and near-shore deposits. Thickness changes across W-plunging folds (Fig. 7) indicate that the folds were active during and after deposition of the transitional member. They likely initiated during deposition of the transitional member and did not persist after deposition of the lower Third Creek Member. The growth folds within this member may be precursors to breakup of the hanging wall of the Bannock detachment system (see below). The presence of some recycled Salt Lake Formation clasts within the transitional member supports this interpretation. Third Creek Member. The Third Creek Member crops out extensively in the Weston Canyon quadrangle west of Rattlesnake Ridge and in the Malad City East and Clifton quadrangles (Figs. 4 and 7). In the Weston Canyon quadrangle, this unit lies without noticeable discordance on the transitional member (Fig. 8). The lower contact is marked by the first persistent conglomerate beds or the base of the upper prominent tuffaceous marker bed where conglomerates are absent. The conglomerate beds are distinctive because they are rich in colored quartzite clasts derived from the Brigham Group. An unfaulted section of Third Creek Member between Rattlesnake Ridge and the Clarkston-Junction fault exposes a minimum thickness of 1.7 km (Fig. 7). The Third Creek Member is thinner in the Malad City East and Clifton quadrangles to the north where 320–1100 m of this unit are preserved (Janecke and Evans, 1999; Carney, 2002). Though Long (2004) does not report any Third Creek Member in the Henderson Creek quadrangle, his 380 m upper conglomerate unit of the Salt Lake Formation (Tscu) is dominated by lithologies that are typical of the Third Creek Member, and these two units may be correlative (Fig. 8). The base of the Third Creek Member contains a 9.6 Ma tephra (Janecke et al., 2003), and the top contains the 4.4 Ma Kilgore ash (M. Perkins, 2005, personal commun.) in the Malad City East and Clifton quadrangles. The absence of the regionally-extensive 6.62 Ma Blacktail Creek and 6.29 Ma Walcott ashes within the upper Third Creek Member in the Malad City East and Clifton quadrangles suggests that there may be an unconformity within the Third Creek Member in those areas (Janecke et al., 2003). This hypothesized missing stratigraphy might be preserved in the unusually thick and continuous Third Creek Member exposed in the Weston Canyon quadrangle (Fig. 8). Our correlation of Long’s (2004) upper “Cache Valley Member” conglomerate unit with the Third Creek Member is supported by the similar 9.6 Ma age of tephras at the base of these two conglomerate units. Within the Weston Canyon quadrangle, the Third Creek Member consists mostly of gray to tan or white medium- to coarse-grained calcareous and tuffaceous sandstone, tuffaceous and/or calcareous siltstone, white to tan nonpetroliferous

limestone, poorly consolidated silvery-gray primary tephra, and lesser pebble to cobble conglomerate. Tuffaceous and limestone lithologies account for ~80%–90% of the Third Creek Member but are difficult to study because they are less consolidated and more poorly exposed than conglomerate beds. Conglomerate beds are 2–8 m thick, structureless to poorly bedded, well- to moderately-sorted, with subangular to rounded pebbles and cobbles (Fig. 10B). Locally, several normally-graded beds are present within a larger conglomerate package. White to gray, poorly consolidated limestone, calcareous sandstone, tuffaceous siltstone, or tephra usually overlie conglomerate beds. The conglomerates have sharp erosional bases or 10–100 cm gradational contacts with underlying gray, tan, or white tuffaceous and calcareous sandstone, pebble-granule sandstone, limestone, and/ or tephra. Individual conglomerate beds are discontinuous, may be lens shaped in map view, and crop out through along-strike distances of 0.2–2.2 km (Figs. 4 and 7). Locally, lateral accretion structures are present in smaller lensoid conglomerate beds. Two observations of clast imbrication suggest an east-directed paleoflow direction. East or west paleoflow is consistent with the observed north-south pinch-outs of the conglomerate beds. To the northwest in the Malad City East quadrangle, a westwardflowing conglomeratic Gilbert-type fan delta in the Third Creek Member was shed from a rising paleo-Oxford Ridge (Janecke and Evans, 1999). Clasts within the Third Creek Member are composed of red, purple, and off-white Brigham Group quartzite, gray Paleozoic limestone and dolomite, bright white Paleozoic quartzite, and recycled tuffaceous siltstone and limestone clasts from the underlying Salt Lake Formation (Figs. 9 and 10B). In the Weston Canyon quadrangle, 46% of clasts are Brigham Group, 35% are Paleozoic, and 19% are recycled green to tan tuffaceous siltstone and limestone from the Salt Lake Formation (n = 21) (Fig. 9), but compositions vary widely from bed to bed. Some conglomerate beds consist of ≥80% of the recycled clasts. Recycled clasts within conglomerates of the Third Creek Member in the central and western part of the Weston Canyon quadrangle are dominated by green to light green tuffaceous siltstone. However, in the Deep Creek half graben on the western edge of the quadrangle, recycled Salt Lake clasts are dominantly derived from tan lacustrine limestone. Interpretation and Unroofing. Clast count data from Salt Lake Formation conglomerates in the Malad City East and Clifton quadrangles reveal an unroofing sequence of middle Paleozoic (early) to Neoproterozoic rocks (late) (Janecke and Evans, 1999; Janecke et al., 2003). The addition of 27 new clast counts from this study and 52 counts from Long (2004) expand the evidence for sequential unroofing stratigraphically lower into the Skyline Member and the Wasatch Formation (Figs. 6 and 9). Upper and middle Paleozoic rocks were eroded during deposition of the Wasatch Formation. Fault-induced erosional exhumation eventually exposed upper Neoproterozoic Brigham Group rocks during deposition of the Third Creek Member (Figs. 6 and 9). The Pocatello Formation in the footwall of the detachment

Evolution of a late Cenozoic supradetachment basin fault was never exposed and did not supply sediment to the Salt Lake Formation (Janecke et al., 2003). A sudden influx of clasts recycled from the older Salt Lake Formation occurred during deposition of the Third Creek Member (Fig. 6). The Third Creek Member differs from the underlying laterally continuous Cache Valley Member by preserving a wide range of depositional environments in aerially restricted subbasins. Overall, this unit records near-shore lacustrine deposition with episodic input of coarse clastics from uplifting highlands. Depositional environments in the Third Creek Member include near-shore freshwater to brackish (?) lacustrine, fluvial-deltaic, fluvial, and beach environments. The influx of recycled Salt Lake clasts in this unit (19%) suggests that some parts of the previously intact supradetachment basin were uplifted and eroded. These distinctive recycled clasts record hanging wall breakup and basin reorganization above the Bannock detachment system (Fig. 11). In the southwest portion of the depositional basin Cache Valley lithofacies were still being deposited when Third Creek lithofacies were being deposited to the north and east. The coarsest facies in the Third Creek Member are localized in the north in the Malad City East and Clifton quadrangles. These are also the only areas that preserve pebble to cobble conglomerate of the overlying New Canyon Member (Janecke and Evans, 1999; Carney et al., 2004). This southward fining trend in the Salt Lake Formation continues into northern Utah into the Clarkston Mountain and Junction Hills areas (Goessel et al., 1999; Oaks, 2000; Biek et al., 2003) and may be due to concentrated extension and uplift near Oxford Ridge. Summary. New work in the Weston Canyon and Henderson Creek quadrangles shows that small amounts of extension began during Eocene time. Lacustrine facies of the Cache Valley Member vary and reflect increasing water depths adjacent to emergent intrabasinal fault blocks. This member is widespread and fairly uniform in thickness in the Malad Range. New tephra-correlation dates from the Salt Lake Formation within the Henderson Creek quadrangle provide better age control within the Skyline and Cache Valley members, show that deposition of the Skyline Member was under way by late middle Miocene time, and confirms that deposition of the Cache Valley Member began just before 10.21 ± 0.03 Ma (Long, 2004). Another important finding is that the Third Creek Member is up to four to five times as thick in the Weston Canyon quadrangle (up to ~1700 m thick) as it is in the Malad City East and Clifton quadrangles. No conglomeratic New Canyon Member lithofacies were identified in the Weston Canyon quadrangle despite the great unfaulted thickness of the Salt Lake Formation. This may reflect southward fining within the upper Salt Lake Formation and suggests that New Canyon lithofacies change laterally into Third Creek lithofacies. STRUCTURAL GEOLOGY The Cenozoic structural geology of the Bannock and Malad ranges records multiple episodes of three-dimensional extension. Consequently, spatial patterns of faults and folds are complex

185

and overlapping. The oldest structures are E-striking normal to oblique-slip faults of Eocene age localized in the northern Henderson Creek quadrangle (Fig. 4). The faults were coeval with the final stages of Sevier-aged eastward shortening in the thrust belt, but developed in the passively transported older thrust sheets riding “piggyback” above the younger, deeper thrusts. Several episodes of ENE-WSW to E-W extension initiated in latest middle Miocene time. The oldest normal faults dip west and localized deposition of the basal Skyline and Collinston members of the Salt Lake Formation (Goessel et al., 1999; Janecke et al., 2003; Long, 2004). Small, conglomerate-filled half grabens began forming before ca. 11–12 Ma and ended before 10.21 ± 0.03 Ma and are localized in a fairly narrow N-S–trending belt (Goessel et al., 1999; Long, 2004). Starting shortly before 10.21 ± 0.03 Ma, the Bannock detachment system developed and became the dominant structure in the area. Its complex evolution from a brief yet extensive early translation phase to a later protracted breakup phase involved the creation of an isostatic anticline (Oxford Ridge anticline), propagation of a new breakaway fault at low angles, and excision and internal extension within the hanging wall of the detachment system (Janecke et al., 2003; Carney and Janecke, 2005; this study). Roughly 50% extension occurred during this main phase of extension (Carney, 2002). Many of the extensional folds in this area date from the detachment phase of deformation. A poorly-characterized episode of cross faulting disrupted the detachment faulting locally, but the main cross-cutting phase of deformation began with Basin and Range–style faulting in middle Pliocene time (Janecke et al., 2003; Carney and Janecke 2005). From west to east, the younger faults include the northern Wasatch and Deep Creek faults and the Clarkston-Junction and Dayton-Oxford strands of the West Cache fault zone (Fig. 1). These faults account for ~10% extension and exposed the footwall of the Bannock detachment fault for the first time within postdetachment horst blocks like the Clifton and Dry Creek horsts (Fig. 4) (Janecke and Evans, 1999; Carney and Janecke, 2005). Bannock Detachment System This field trip focuses on the Bannock detachment fault system, and its most prominent fault, the Clifton detachment. Several low-angle normal faults comprise the Bannock detachment system. The structurally highest Clifton fault is the most continuous and omits the most section across it, but in the Clifton horst there are at least three smaller, offset, low-angle normal faults in its footwall (Fig. 4) (Carney and Janecke, 2005). The Clifton detachment strikes NNW and persists >39 km N-S. It has a low-angle geometry through its entire extent, with an average dip of 15° to the WSW, and exhibits slight waviness along strike. Sparse but excellent exposures of the fault surface (Fig. 10C) confirm its low dip (strike of 158.6° and dip of 19.6° WSW ± 11.6°; n = 5) and WSW slip direction. Slickenlines trend 255° and plunge 33° on average; n = 9 (Fig. 12D). These exposures and breccia bodies coincide with the flat-on-flat

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Data from this field guide (n=43)

Equal Area

Combined fold axis (n=78) Data from Carney and Janecke (2005; n=35)

e irdl y ag dat Carne e k from Janec d 5 n a 200 ) (

1

combined data girdle

1

1

Bannock detachment fault (n=5) Black indicates mean vector Average strike and dip 158.6; 19.6 SW +/- 11.6 Pocatello Foramtion (n=16) Black indicates mean vector Average strike and dip 156.0; 19.2 SW +/- 4.1 Salt Lake Formation (n=6) Black indicates mean vector Average strike and dip 145.0; 29.1 SW +/- 14.2

Equal Area

B

Southern Rattlesnake Ridge flat-on-flat

ge era av

38, 281

-d NW

l ing ipp

im

b

Oxford Ridge anticline

ave ra fau ge de lt (1 t ac 57; hm 20 ent SW )

A

plunge and trend of fold axis 0, 164 1, 160 1, 155

n=78

SW-dipping fold limbs 331-45

plunge and trend of fold axis

av er ag

average pole to fault

33, 255 eS W -d

ip

NW-dipping fold limbs 27-39 pi

ng

lim

plunge and trend of average slickenline

n=9 (slickenlines) n=5 (fault planes)

b

n=17 Equal Area

C

W-plunging growth folds in Salt Lake Formation

Equal Area

D

Bannock detachment fault and slickenlines

Figure 12. Stereograms of important structural features. (A) Measurements documenting the flat-on-flat relationship along southern Rattlesnake Ridge in the Weston Canyon quadrangle. Plot shows bedding in the hanging wall (Salt Lake Formation) and footwall (Pocatello Formation) and compares these with the attitude of the intervening Bannock detachment fault. Note the overlap of the three mean vectors. This is consistent with a flat-on-flat relationship. (B) Oxford Ridge anticline in the Clifton and Weston Canyon quadrangles. Bedding and foliation orientations from the Pocatello Formation in the footwall of the Bannock detachment fault system define a subhoriztonal NNW-SSE–trending anticline. (C) Attitudes of growth folds south of Fivemile canyon within the Transitional and lower Third Creek members of the Salt Lake Formation. Bedding from the Salt Lake Formation defines an overall W-WNW–trending, moderately plunging fold axis. Note the plunge of the fold axis is very similar to the bedding dips. (D) Slickenlines and fault planes from southern Rattlesnake Ridge. Note the overall WSW-trending slickenline measurements and low dip of the Bannock detachment fault. Also note that the trend of the Oxford Ridge anticline (B) is approximately parallel to the strike of the Bannock detachment fault (D).

Evolution of a late Cenozoic supradetachment basin portion of the fault and preclude interpretations of the contact as an unconformity. The Neoproterozoic Pocatello Formation comprises the footwall of the Clifton detachment, whereas hanging wall rocks vary along strike from rocks of the Neoproterozoic to Cambrian Brigham Group, Cambrian to Ordovician Formations, and the Miocene-Pliocene Salt Lake Formation (Fig. 8). In the Weston Canyon and parts of the Clifton quadrangle, the Clifton detachment juxtaposes late Cenozoic Salt Lake Formation with metamorphosed Neoproterozoic Pocatello Formation and omits up to 6 km of rocks. None of the clasts within conglomerates of the Salt Lake Formation above the detachment fault were derived from the Pocatello Formation beneath the detachment (Figs. 6 and 9). The hanging wall of the Clifton detachment contains lowangle and high-angle normal faults. These younger WSW- and ENE-dipping normal faults sole into or are cut by the master Clifton detachment fault. The spacing between hanging wall faults

A (W)

varies widely but is generally close in the Clifton quadrangle north of the lateral ramp near Fivemile Creek (see below). There are no faults in the hanging wall of the detachment south of the lateral ramp (Fig. 4). Flat-on-Flat Geometry across the Detachment Fault Along most of Rattlesnake Ridge, footwall and hanging wall rocks have approximately the same strike and dip and parallel the intervening detachment fault (Fig. 12A). A potentially complicating factor of this analysis is that 3-point determinations of the dip of the detachment fault are 10°–20° lower than the dip measured in outcrop. This difference reflects the position of the 3-point measurements closer to the crest of the Oxford Ridge anticline, whereas direct measurements sampled the limbs of the anticline farther west. Our structural cross sections (Fig. 13) suggest that bedding in the Third Creek Member may also be parallel to the detachment fault north of Fivemile Creek. This

A‘ (E)

Oxford Ridge anticline

1000 m

187

undivided Quaternary deposits

Q

Tslu undivided Salt Lake Formation

Da on

yt

Ttc Third Creek Member of SLF or

xf

-O

Ttc-cv Third Creek-Cache Valley Transitional Member of SLF

d

Tcv Cache Valley Member of SLF t

ult on fa

Clift

Ttc Ttc

Clifton Ttc

Zps

fault

Tcv

Pzu Zps

1000 m

e th lt of fau nd d ra or st xf st n-O Ea yto Da

ul

fa

Ttc

2000 m

n Cliftuolt fa Tcv Ba nn oc kf au lt

Pzu undivided Paleozoic rocks Zp Neoproterozoic Pocatello Formation undivided pre-Cambrian rocks projected dip of bedding Q bedding form line

Tslu

fold axial trace

Zps Clarkston strand of the West Cache fault

Basin-and-Range normal faults

Zps

Clifton fault of the Bannock detachment fault system pre-Cambrian, undiff.

Bannock fault of the Bannock detachment fault system

MSL

B (W)

B‘ (E)

D

ay

to

1000 m

n-

Ttc

O

xf

or

2000 m

d

fa

ul

t

Ttc Tcv

Ttc

Ttc Pzu

Zps

slight stratigraphic thinning is hypothetical

1000 m Ttc Clarkston strand of the West Cache fault

Ttc-cv Clifton fault

Tcv Ttc-cv t l au nf ifto Cl

Ttc Tcv Ttc-cv Ban noc k fa ult

Oxford Ridge anticline

Q Tslu

Zps Zps

Zps pre-Cambrian, undiff.

MSL

Figure 13. Geological cross sections from the Weston Canyon quadrangle (see Fig. 7 for locations). A–A′ crosses the upper flat of the Bannock detachment system and shows the eroded Clifton fault merging with the older and underlying Bannock fault near the crest of the Oxford Ridge anticline. Note the cut-off angle between strata above and below the Clifton fault near the anticline. This geometry is expected during the process of excision. B–B′ crosses the lower flat of the Bannock system and shows similar geometries to the northern section.

A.N. Steely et al.

Hanging wall strata Tcv-m

youngest

oldest

?

12.0

9.0

Upper flat

CZcm

10.0

Tcv-u

11.0

Ttc

(hanging wall and footwall)

Brecciated strata within thin fault blocks of the BDF

Tcv-l

5.0

4.0

1.0

Lower flat

2.0

Tcv-cgl

3.0

0

(hanging wall and footwall) (hanging wall and footwall)

Tcv-m

Cbl

6.0

Tcv-u

7.0

Lateral ramp

Ttc-cv

8.0

CZcm

Along-BDF-strike distance from the southern tip of Rattlesnake Ridge (in km)

188

(Tcv-cgl) Cache Valley Formation-basal with conglomerates (Tcv-l) Cache Valley Formation-low (Tcv-m) Cache Valley Formation-middle (Tcv-u) Cache Valley Formation-upper (Ttc-cv) Third Creek-Cache Valley Transitional unit (Ttc) Third Creek Formation-low (Cbl) Blacksmith Formation (CZcm) Camelback Formation

Figure 14. Strata in the hanging wall of the Bannock detachment fault (BDF) as a function of distance from the southern tip of Rattlesnake Ridge. These are compared against the units within thin fault blocks (horses) of the underlying detachment fault. Note the northward younging of hanging wall strata across the Fivemile lateral ramp. The brecciated Camelback Mountain Formation coincides with the upper and lower footwall and hanging wall flats, whereas the Blacksmith Formation coincides well with the lateral ramp of the detachment fault. See text for additional discussion of the lateral ramp. Although the significance of younger rocks in horses along the ramp is uncertain, it helps to define the extent of the lateral ramp.

flat-on-flat geometry is one of 10 arguments that suggest the Bannock detachment fault formed and slipped at low angles (Carney and Janecke, 2005). The depth to the detachment fault while it was slipping can be constrained by the thickness of the supradetachment basin fill above the fault. In the Weston Canyon quadrangle, only the Salt Lake Formation lies above the detachment fault, and post detachment deposits such as the Quaternary-Tertiary piedmont gravels of Janecke et al. (2003) are only a few hundred meters thick at most and were deposited after the death of the detachment fault. This suggests that unless a significant thickness of unrecognized post–New Canyon Member deposits were eroded from the area, the ~2.5–4.3 km thickness of the Salt Lake Formation comprised the only rocks above the detachment fault. Because of the unique flat-on-flat geometry, this further suggests that the originally subhorizontal Bannock detachment fault was slipping at depths of <2.5–4.3 km. Lateral Ramp in Hanging Wall and Footwall of the Detachment Fault In contrast to areas areas north and south, hanging wall and footwall rocks are truncated by the detachment fault in the vicinity of Fivemile Creek (Fig. 7). Figure 14 illustrates changes in hanging wall stratigraphy as a function of distance along-strike of the detachment fault. Along the southern segment of the fault, the lowest stratigraphic levels of the Salt Lake Formation within the Weston Canyon quadrangle are faulted against the Pocatello Formation. However, starting ~4 km north of the southern tip of Rattlesnake Ridge, progressively younger rocks are faulted against the Bannock detachment fault (Fig. 14). At the sharp northward bend in the detachment fault along Fivemile Creek, the fault has cut significantly upsection and faults lower Third Creek Member against the Pocatello Formation. This upsection truncation of hanging wall units is expressed in map view near Fivemile Creek (Fig. 7). Mapping of the Pocatello Formation in the footwall shows a complimentary pattern of stratal truncations. We suggest that stratal truncation of footwall and hanging wall rocks is created by footwall and hanging wall lateral ramps in the Bannock detachment fault, herein called the Fivemile lateral ramp. The flat-on-flat geometry north and south of the lateral ramp (Fig. 14) and the thickness of the Tertiary units truncated by the ramp define the original height of the lateral ramp. We estimate an original height of ~850–1100 m along a fault-strike distance of ~4–4.5 km. This estimate takes into account the full thickness of the transitional member (198–368 m), most of the measured minimum thickness of the Cache Valley Member (~600 m), and a small thickness of the Third Creek Member. The amount of truncated Third Creek Member is poorly constrained, but an estimate of 50–150 m is reasonable based on map patterns. Overall, these data show that the Fivemile lateral ramp had an original south-southeast dip of 13.0° ± 2.3° and was approximately perpendicular to the strike of the detachment fault. The lateral ramp is spatially coincident with a change in the lithology of fault-bound slivers along the detachment fault (Fig. 14).

Evolution of a late Cenozoic supradetachment basin

42˚15´

Dayton

fault East Cache

Valley

ult

che fau lt

Cache

s fa Hill ville

lls We

41˚45´

tch

s

ain

sa Wa

unt

Mo lt

fau

0

Basin-and-Range normal faults (dotted where concealed) 25 KM

Logan

e Bear River Rang

ion nct -Ju ton rks

Bedrock

42˚

Lewiston fault

fault -Oxford

Cla

CM

Preston

East Ca

HF

Ridge Oxford

Area of Fig. 4

Samaria Mtn West Hills

111˚ 45’

112˚

112˚ 15’

e Rang Malad ult ch fa asat

Excision of Hanging Wall Rocks Cross sections across the Bannock system in the Weston Canyon quadrangle (Fig. 13) suggest that two detachment faults are present east of the crest of the Oxford Ridge anticline. The lower fault (Bannock fault) is preserved in two small outcrops near Fivemile quarry and one ~2 km south of the quarry (Fig. 7). In all three locales hanging wall rocks consist of Cache Valley Member and bedding is parallel to subparallel with the dominantly east-dipping detachment fault (Fig. 13). The higher detachment fault strand is inferred above the current level of exposure (Fig. 13). Farther east it may be faulted below Cache Valley on the east strand of the Dayton-Oxford fault zone. The higher strand merges with the Bannock fault near the crest of the Oxford Ridge anticline. East of the Oxford Ridge anticline, the higher strand cuts down section through the Salt Lake Formation to the west (Fig. 13). Fault excision is an important process of the Bannock detachment system in the Clifton quadrangle to the north (Carney and Janecke, 2005). We believe that excision also played a major role in the development of the Bannock system in the Weston Canyon quadrangle and is responsible for the Clifton detachment fault cutting upsection into the synextensional Salt Lake Formation (Fig. 16). A minimum of ~4 km of slip is documented on the Clifton fault in the Clifton quadrangle

112˚30´

W ley Val lad Ma

Major changes in the amount of hanging wall deformation occur across the Fivemile lateral ramp. Above the southern, lower flat along Rattlesnake Ridge, the hanging wall is unfaulted, whereas north of the Fivemile lateral ramp the hanging wall is internally faulted above the northern upper flat (Figs. 4 and 7). Because the lateral ramp cuts out at least ~850–1100 m of stratigraphy, the depth to the northern flat during transport was <1.6–2.6 km, depending on whether the additional thickness of the New Canyon Member was once present or not. We suggest that the increased faulting above the northern, upper flat is a result of decreased overburden thickness above the detachment fault. This pattern is consistent with two-dimensional structural models of detachment faults that show more highly extended thin hanging walls and less faulted thicker hanging walls. Detailed gravity data confirm the south to north decrease in basin depth across the lateral ramp that our model predicts (Fig. 15). The Fivemile lateral ramp is an ENE-trending structure that coincides with a major north to south change in structural style above the Bannock detachment fault. The westward projection of the ramp coincides with (1) the interaction zone between the younger Deep Creek fault and Clarkston-Junction strand of the West Cache fault; (2) an anomalous NE-striking part of the Clarkston-Junction fault; (3) the ENE-trending fault block at the south end of the Skyline half graben; (4) a major segment boundary in the Wasatch fault; and (5) a bedrock ridge beneath Malad Valley (Figs. 4 and 15). Eastward the Dayton-Oxford fault has a branch point and a left bend along this trend. The alignment of these major lateral structural changes might reflect structural inheritance within this ENE-trending belt (Fig. 4).

189

Brigham City

41˚30´

Figure 15. Isostatic residual gravity contour map of north-central Utah and southeastern Idaho. Contour interval is 4 mGal. Hachures indicate gravity lows. Most of the anomalies reflect the density contrast between Cenozoic basin fill and pre-Cenozoic bedrock. Note that many of the Basin and Range normal faults are associated with prominent gravity gradients, with lower values on the downthrown side of the fault. New detailed data sets include Oaks et al. (2005), Eversaul (2004), Goessel (1999), and Smith (1997). HF—Hawk fault; CM—Clarkston Mountain.

where other data further suggest a total displacement of ~15 km (Carney and Janecke, 2005). Folds Within the Weston Canyon quadrangle, there are extensional folds of several orientations. Descriptions of these structures will be limited to those that occur in the hanging wall and footwall of the Bannock detachment system (Fig. 4). The folds within the study area can be subdivided into two main geometries: NNNW–trending longitudinal folds parallel to the normal faults of the Bannock detachment system and transverse folds with W-WSW trends perpendicular and/or oblique to the Bannock detachment system.

A.N. Steely et al.

190

A) Ttc

Ttc Tcv

ramp-on-flat

future Clifton fault Ttc-cv Tcv Bannock fault

fu

B)

tu

Ttc

re yt Da

Clifton fault Ttc

on

Ttc-cv

-O

Ttc -flat flat-on

xfo rd u fa

Tcv Bannock fault (inactive)

lt

Figure 16. Schematic illustration of excision of the Bannock detachment fault system. (A) Unloading above the Bannock fault creates a broad isostatic uplift (Oxford Ridge anticline) that back-tilts the detachment fault into an unfavorable geometry for continued slip. (B) The Clifton fault cuts upsection into the hanging wall (excises) and merges with the Bannock fault at the crest of the anticline. Flat-onflat geometries suggest that slip on the Clifton fault has removed the ramp-flat relationship that would initially develop above the excising fault and transported it down-dip. This is schematically shown with the location of stars above the Clifton fault. Note the expected cut-off between strata above and below the Clifton fault near the anticline. Dashed red lines indicate future faults of the Dayton-Oxford fault. Legend is same as Figure 13.

Longitudinal Folds One longitudinal fold deforms rocks in the Weston Canyon quadrangle. This fold continues northward into the Clifton quadrangle where Carney and Janecke (2005) first described the structure and named it the Oxford Ridge anticline. In the Weston Canyon quadrangle the axial trace of the Oxford Ridge anticline trends along Rattlesnake Ridge for ~9 km before plunging south beneath Quaternary deposits at the southern end of the ridge (Fig. 7). Regionally, the Oxford Ridge anticline is more than 40 km long (Carney and Janecke, 2005), similar to the >39 km exposed length of the Bannock detachment fault The Oxford Ridge anticline is defined by changes in the dip of the Bannock detachment fault, the Neoproterozoic Pocatello Formation in the footwall, and Salt Lake Formation in the hanging wall (Fig. 7). Within the Weston Canyon quadrangle, 43 footwall bedding and foliation measurements define the Oxford Ridge anticline as an open fold trending 155° and plunging 1° southeast (Fig. 12B). This is in close agreement with the 345° trend and 0° plunge of the Oxford Ridge anticline in the Clifton quadrangle to the north (Carney and Janecke, 2005). Combining these two data sets yields an average 160° trend and 02° plunge (Fig. 12b; n = 78). A ~7 km long N-S–trending anticline in the SE part of the Henderson Creek quadrangle (Fig. 4) appears to be a composite double rollover anticline. Its west-dipping limb formed during slip on the NE-dipping Steel Canyon fault between ca. 9.6 and

<9.2 Ma. The east-dipping limb dates from the post–4.4 Ma Basin and Range phase of faulting on the west-dipping Deep Creek fault (Long, 2004). Transverse Folds Folds with WNW-SW trends in the Weston Canyon quadrangle can be further subdivided into those that deform both footwall and hanging wall rocks of the Bannock system and those that deform only hanging wall rocks. Folds that deform the entire Bannock detachment system are located near Fivemile Creek west of the Dayton-Oxford fault (Fig. 7). These folds plunge WSW and have tight to open geometries. Cross-cutting relationships show that these are younger than west-plunging folds 0.7 km farther south that deform only hanging wall rocks. Folds that only deform hanging wall rocks are localized primarily in the upper Cache Valley to lower Third Creek members of the Salt Lake Formation south of Fivemile Creek (Fig. 7). The northern three folds also define an overall syncline similar in scale to a prominent syncline located ~2 km to the south. Mapping and stereonet analyses of these structures define an average trend of 281° and plunge of 38° (Fig. 12C) with wavelengths of ~1.3–1.6 km for the larger-scale, more widely spaced folds. Data discussed in the “Stratigraphy” section show that the transitional unit thickens and thins across these folds. Interpretation The longitudinal Oxford Ridge anticline was interpreted as an isostatic fold that formed in response to crustal unloading above the Bannock detachment fault (Carney and Janecke, 2005). Our new data strengthen this interpretation by showing that the fold is perpendicular to slip on the detachment fault (Fig. 12) and by documenting that the Oxford Ridge anticline deforms the entire N-S extent of the detachment fault. The formation of an isostatic anticline suggests significant amounts of extension on the Bannock fault. The current westward dip of both bedding and fold axes west of Rattlesnake Ridge is due to a combination of steepening on the west flank of the Oxford Ridge anticline and slip on the E-ESE– dipping Clarkston-Junction strand of the West Cache fault. Two generations of subparallel folds plunge west. The younger set near Fivemile Creek deforms the lateral ramp in the Bannock detachment fault and accentuates the map view truncation of hanging wall rocks. We interpret these folds as forming after deposition of the Third Creek Member, likely after main-stage movement on the Bannock detachment system. The older W-plunging folds south of Fivemile Creek are confined to the hanging wall and are interpreted as growth folds active during deposition of the Salt Lake Formation. Folding likely began during deposition of the lowermost transitional member, or less likely, during deposition of the uppermost Cache Valley Member. Basin and Range Normal Faulting Steep to moderately dipping Basin and Range normal faults generally strike north and dip both east and west, forming horsts,

Evolution of a late Cenozoic supradetachment basin grabens, and half grabens (Figs. 1 and 4). Prior studies (Machette et al., 1992; Evans and Oaks, 1996) have noted the major, active, west-dipping Wasatch and East Cache fault zone. Our mapping indicates that the intervening Deep Creek and West Cache fault zones are also major structures at this latitude. The Clarkston–Junction Hills strand of the West Cache fault zone dies out northward near 42°7.5′ but overlaps for more than 30 km with the DaytonOxford normal fault to the east. Between these two east-dipping en echelon strands of the West Cache fault zone, the Salt Lake Formation and the Bannock detachment system dip westward. The west-dipping Deep Creek normal fault bounds intermontane basins in the Malad and SE Bannock ranges. It has the greatest throw in the Clifton quadrangle, where a ~1–1.5-km-deep sedimentary basin is preserved in its hanging wall (Evans et al., 2000; Eversaul, 2004). Southward, the Deep Creek fault becomes more complex, steps right in several relay ramps and breached relay ramps, and loses displacement. A slip minimum coincides with an E-W–trending bedrock ridge of Ordovician rocks in the hanging wall of the Deep Creek fault in the northeastern Henderson Creek quadrangle and NW Weston Canyon quadrangle (Fig. 4). The location of this slip minimum is consistent with the location of an interaction zone between the Deep Creek and West Cache fault zones north of the bedrock ridge of Ordovician rocks (Fig. 4). Numerous small to moderate displacement normal faults in this interaction zone cross Weston Canyon and connect the oppositely dipping West Cache and Deep Creek faults across the Dry Creek horst block (Fig. 4). This complex fault network probably transfers extension between the two normal faults and explains both the drop in displacement across the Deep Creek fault that occurs south of the interaction zone and the drop in slip on the West Cache fault north of the interaction zone. Wasatch Fault and Relay Ramp The Wasatch fault is an active, north-striking, west-dipping normal fault system that extends from central Utah to southeast Idaho (Machette et al., 1992). In the Henderson Creek quadrangle, the Wasatch fault system consists of two right-stepping N-NW– striking segments, which are separated by a relay ramp and bedrock ridge in the hanging wall (Figs. 1 and 4). The Clarkston Mountain segment is a 19-km-long segment along the southwestern edge of the Malad Range and Clarkston Mountain (Machette et al., 1992). Throw on this segment decreases to the north, as it is progressively transferred to the Malad City segment. The Malad City segment of the Wasatch fault defines the western boundary of the northern part of the Malad Range (Fig. 4). The two faults overlap for ~10 km, and form a 2.5–3.5-km-wide en echelon right step. New gravity data collected in the greater Cache Valley region highlight the relative throws of the Basin and Range normal faults, locations of subbasins and buried structural ridges (Fig. 15) (Oaks et al., 2005). The aerial extent and amplitude of the gravity lows are a proxy for the size and depth of the Cenozoic basins. Using this as a guide, two basin systems dominate the gravity field. The Cache Valley basin system is in general

191

deeper and wider than the basins west of the Wasatch fault. However, in the central part of the area, between 42°7′ and 41°45′, the two-basin system is replaced by three, elongate, subparallel NNW-trending basins. The change from two to three basins in this latitudinal band shows that the east-dipping West Cache fault zone is a major structure in this region and rivals the Wasatch and East Cache fault zones in its throw. FIELD TRIP This field trip guide starts at the Woodruff/Samaria exit of I-15, 2.1 mi north of the Utah-Idaho state line and 95.2 mi north of Exit 311 of I-15 (400 S.) in Salt Lake City, Utah. Mileages are given from the north end of the Woodruff/Samaria off-ramp. All UTM coordinates are in reference to Zone 12 of the NAD27 datum. Mileage Interval (Cumulative) Description 0.1

0.1

Turn right on Woodruff Lane at the end of the Woodruff/Samaria off ramp and park in pullout by 2000 W.

Stop 1: Overview of Southern Malad Range; UTM 0399932E 465422N The purpose of this stop is to discuss recent work by Long (2004) and Long et al. (2004) in the Henderson Creek quadrangle. This work identified the probable location of the monocline at the west edge of the Cache-Pocatello culmination and refined the age, sedimentological, and structural interpretations of the Skyline and Cache Valley members of the Salt Lake Formation. Work in the Cache Valley Member characterized two different lacustrine facies and documented the Late Miocene age of the Steel Canyon normal fault. Mileage Interval (Cumulative) Description 0.2

0.3

Return to Woodruff Lane and drive west ~0.2 mi to the intersection with Old Highway 191. Turn right (north). 8.9 9.2 Drive north to 1500 S. and turn right (east). 3.7 13.2 Drive east 0.8 mi on 1500 S. and cross under I-15 where 1500 S. becomes Two Mile Creek road. Cross the Wasatch fault into the Malad Range. Continue 2.9 mi up Two Mile Creek to Stop 2. Alternate route (if roads are very wet and muddy—skip stop 2): Do not turn right at 1500 S. from Old Hwy 191—continue straight (north) into Malad City ~1.5 mi. Turn right and get on I-15 North. Drive north and exit at Hwy 36 East (Exit 17). Turn right (east) and drive on Hwy 36 for 7 mi past Deep Creek Reservoir to a large pull-out on the right. Rejoin the field guide at Stop 3 and reset mileage.

A.N. Steely et al.

192 Stop 2: Examine Altered Tuffaceous Lake Beds; UTM 0403322E 4668926N

Make a brief stop along Two Mile Canyon to examine green altered tuffaceous siltstone of the Cache Valley Member of the Salt Lake Formation. Notice that the tuffaceous sedimentary rocks are structureless to finely laminated. Other sedimentary structures are lacking. These deposits and interbedded lacustrine limestones up and down section suggest that ash fall beds were deposited below wave base in an open lake. The Cache Valley Member in this area overlies ~200–250 m of the conglomeratic Skyline Member of the Salt Lake Formation. As we drive east along this road we will cross a narrow horst block of Cambrian carbonates shortly before a T intersection with Highway 36. The west-dipping Red Knoll fault on the western margin of this horst block was the basin-bounding normal fault during deposition of the Skyline Member (Janecke et al., 2003; Long, 2004). The Red Knoll fault was inactive when the overlying Cache Valley Member was deposited across the older fault block. Mileage Interval (Cumulative) Description 0.7

(13.9)

2.7

(16.6)

1.7

(18.3)

Skyline road joins from the right; continue east (left). Continue on Two Mile Road over the Malad Range divide down to a T-intersection with Hwy 36. Turn left (north) onto Hwy 36 and drive 1.7 mi to a large pull-out on the left (west) side of the road just north of milepost 108.

Stop 3: Extensional Monocline and Breakup Deposits of the Salt Lake Formation; UTM 0406959E 4669904N This stop will examine facies within the Third Creek Member of the Salt Lake Formation and observe extensional monoclines in the hanging wall of a concealed normal fault. Cross to the east side of Hwy 36 and locate a gate in the fence near the obvious wash. This stop will illustrate (1) the variable lithologies of the Third Creek Member; (2) the presence of voluminous clasts of reworked tuffaceous sediment from the underlying Cache Valley Member of the Salt Lake Formation; and (3) some of the well-developed extensional folds that deform the hanging wall of the Bannock detachment fault. An exposure of pebble conglomerate near road level is composed of recycled clasts of the Cache Valley Member of the Salt Lake Formation. Notice the degree of sorting and rounding of these clasts. This bed and one near the parked cars dip gently down the road (to the NW). Go through the gate and walk up the wash. Notice that the strike of the beds has change to NW and dips steepen as we walk NE through the wash. In beds SE of the wash the Third Creek beds are nearly vertical (up to 87°). The gentle beds along the highway and these steeper beds define

a synformal monoclinal hinge of a NNW-plunging fold (best-fit plunge and trend of the cylindrical fold axis is 19°, 331°; n = 32). Steep beds flatten at structurally higher levels and represent the antiformal hinge of the NW-plunging monocline. Individual tilt panels range from NW- to SW- to NE-dipping and define at least four monoclinal hinge zones with NNW plunges. Altogether, the monoclinal folds define a plunging kink-like rollover fold in the hanging wall of the NE-dipping Hawk normal fault (Evans et al., 2000). The Hawk fault is nowhere exposed and/or is covered by alluvium, but it is needed to explain the stratigraphic and structural relationships on the geologic map. Gravity data across this normal fault shows a ~7 mGal drop from the footwall (SW) to the hanging wall side of the fault, consistent with moderate displacement across this fault (Evans et al., 2000; Eversaul, 2004). The NW-striking Hawk fault formed during Miocene to Pliocene detachment faulting, is antithetic to the Bannock detachment fault, and may be one of the late stage breakup faults that formed above the detachment system during deposition of the Third Creek and New Canyon members of the Salt Lake Formation. Notice exposures of conglomerate on either side of the wash as you walk to the east. Some pebbly beds contain ooids and gastropods in the matrix. We interpret the beds here as the distal toes of alluvial fans at the margin of a lake. Keep walking up the wash to observe an exposure of a thick vitric ash on the left (north) bank of the wash beneath the lake-margin deposits. Tephra correlations show that this ash is the 7.9 ± 0.5 Ma Rush Valley Ash (Janecke et al., 2003). After examining the late Miocene ash, climb up on the slopes SE of the wash to exposures of nearly vertical conglomerate beds. This vantage point provides a better view of the monoclinal fold train in the hanging wall of the Hawk fault. A horst block of Cambrian limestone with a thin carapace of red weathering Eocene Wasatch conglomerates is to the east. This horst block is bounded by N-S–striking normal faults that cut across and postdate the NW-striking faults and folds that we just examined. Cross-cutting relationships like these allow us to distinguish between the many phases of normal faulting and folding in the area. The N-striking normal faults are from the youngest Basin and Range phase of extension. Return to the vehicles, turn around, and drive south on Hwy 36. Mileage Interval (Cumulative) Description 7.0

(25.3)

5.9

(31.2)

0.7

(31.9)

Cross Clarkston-Junction strand of the West Cache fault. Drive along Weston Creek through tilted Salt Lake Formation overlain by Lake Bonneville nearshore and lake deposits. Cross the buried trace of the DaytonOxford fault and also note the thick roadcut of Bonneville-age gravel deposits north of the road. Turn north (left) onto Franklin County D1/Hwy 36 and drive toward Dayton,

Evolution of a late Cenozoic supradetachment basin

4.9

5.4

6.3

(36.8)

(42.2)

(48.5)

Idaho. Notice Rattlesnake Ridge to your left (west side of road) in the footwall of the east-dipping Dayton-Oxford fault. Enter Dayton. At the T-intersection turn left to stay on Franklin County D1 and continue north to Clifton. Turn left on 100 S. at the dark-colored sign for Clifton Cemetery (on right side) and continue on this winding dirt road. Turn left on Cemetery Rd. Drive up in the Bannock Range to the end of the road in Davis Basin. High clearance is advisable. Park at the end of the road and walk NNE up a trail ~3/4 mi to a trail junction between the trail along Oxford Ridge and the E-W trail to Mine Hollow. Leave the official trail and hike SW along a ridgeline until you have a clear view of the rocks to the north. There should be brecciated quartzite underfoot.

Stop 4: View Excision along the Clifton Detachment Fault; UTM 0412275E 4673135N We will stop at a high point with a view to the north of Oxford Ridge. The Clifton detachment fault forms a near dip slope on the WSW edge of Oxford Ridge and we can view it a few kilometers to the north of us. There the fault dips gently WSW and truncates (excises) numerous SW- and some NE-dipping normal faults in its hanging wall (Carney and Janecke, 2005). These hanging wall faults repeat lower Paleozoic rocks, overlying Wasatch Formation, and the Cache Valley Member of the Salt Lake Formation. The rusty-weathering rocks with a dense cover of small trees are the Neoproterozoic Pocatello Formation. Its foliation (and bedding) is subparallel to the detachment in this area. A cross section, detailed description of these relationships, and the evidence for excision are presented in Carney and Janecke (2005). These relationships indicate slip on the Clifton fault at a low dip angle within a few kilometers of Earth’s surface. A short distance NE of here in northern Cache Valley one can see large ancient meanders produced by the outflow stream of Lake Provo. These meander bends are preserved in Oxford Slough and show evidence for northward flow when they were active. Flow directions have reversed since the Bonneville lake system drained, and Oxford Slough is now integrated into the south-flowing Bear River system. Return to vehicles and retrace path back to Franklin County D1 in Clifton. Mileage Interval (Cumulative) Description 6.3

(54.8)

5.0

(59.8)

Turn right (south) on Franklin County D1 and continue to the north end of Dayton. Turn right (west) onto the access road to Fivemile Creek.

1.6

(61.4)

193

Just past the end of the pavement, turn right (north) onto a dirt track. Wind around a small bluff to enter Fivemile quarry.

Stop 5: Faulted Rocks above the Detachment and Conglomerates of the Cache Valley Member; UTM 0416229E 4662973N This stop is located between two east-dipping strands of the Dayton-Oxford fault, which at this latitude is the major basinbounding structure of Cache Valley. Gravity data (Fig. 15) show that the eastern Dayton-Oxford fault has much more throw than the western Dayton-Oxford fault at this latitude, consistent with our structural analysis (Fig. 13). This stop starts in the lower Cache Valley Member of the Salt Lake Formation, located just above a tilted E-dipping portion of the Bannock detachment fault, and requires a short walk to examine thin conglomerate beds within the lower Cache Valley Member. The quarry is developed in very fractured, faulted, and locally brecciated, tuffaceous to non-tuffaceous siltstone and mudstone. The fractures and faults obscure original bedding and are likely a result of this outcrop’s proximity to the underlying detachment fault. Brecciated Brigham Group Quartzite in fault blocks within the detachment fault crop out just west of the quarry (Figs. 7 and 13). The approximately concordant east dips of the Bannock detachment fault and the overlying Cache Valley Member suggests a flat-on-flat geometry. This E-dipping flat-on-flat geometry east of the crest of the Oxford Ridge anticline helps to constrain the evolution of the Bannock detachment system (Fig. 16). To reach the conglomerates, walk NE around the edge of the quarry through a wash cut in green to white tuffaceous siltstone. Tuffaceous beds and limestone account for ~90% of the Cache Valley Member. Walk NE up the slope to the top of the ridge where conglomerates crop out at UTM 0416486E 4663294N. These conglomerates are rare in the Cache Valley Member, only crop out in the lower few hundred meters of section, and are locally present for ~7 km along strike to the south. Angular to rounded clasts within the conglomerates are derived from Paleozoic rocks. Matrix within the conglomerates ranges from calcareous sand to micrite. Micrite-matrix conglomerates are surprising due to the overwhelming dominance of tuffaceous siltstone lithologies and the lack of significant limestone in this section. We interpret these relationships to suggest conglomerate progradation and limestone deposition during times of reduced ash input to the lacustrine environment and surrounding landscape. We also suggest that highlands must be present nearby, likely to the east, to supply the conglomerates during these time periods. These highlands may have formed by an intrabasinal fault. Retrace your route to the main dirt road along Fivemile Creek. Mileage Interval (Cumulative) Description 0.6

(62)

Turn right (west) onto the main dirt road and drive west 0.4 mi.

A.N. Steely et al.

194 0.2

(62.2)

1.0

(63.2)

Turn left to stay on Fivemile Creek road. Drive into Fivemile canyon. At 0.5 mi cross the western strand of the Dayton-Oxford fault onto Neoproterozoic Pocatello Formation. Continue up the canyon through the Scout Mountain Member of the Pocatello Formation. Just west of the main gorge of Fivemile canyon, cross the W-dipping Clifton strand of the Bannock detachment fault and park along the north shoulder of the road.

Stop 6: Lateral Ramp in the Bannock Detachment System; UTM 0414359E 4662120N Walk to the north side of the road and climb the blue-gray outcrop of Cambrian Blacksmith Formation limestone within fault blocks along the detachment fault. From this vantage point, we can see the W- to S-dipping Bannock detachment fault. The purpose of this stop is to discuss evidence for (1) a folded lateral ramp in the Bannock detachment fault system, (2) the Oxford Ridge anticline, and (3) the transitional unit between the Third Creek and Cache Valley members. The folded lateral ramp of the detachment fault can be seen in map view as northward and southward truncation of both footwall and hanging wall rocks against the detachment. To the south along Rattlesnake Ridge, the lowest outcrops of the Cache Valley Member are cut by the fault. There the hanging wall, footwall, and fault are approximately parallel and we interpret this as a flat-on-flat geometry. Here along Fivemile Creek, the fault has cut upsection into the transitional member of the Salt Lake Formation. Northwest of here, near Stop 7, the detachment fault cuts lower Third Creek Member. Our structural cross sections (Fig. 13) suggest that a flat-on-flat geometry may also exist where the Third Creek Member is cut by the fault to the north. Figure 14 shows the N-S changes in hanging wall stratigraphy above the detachment fault. We use these data to interpret an ~850–1100m-high lateral ramp in the detachment fault. Gravity data shows a gravity high along the NNW-trending Oxford Ridge anticline but also provides evidence for the ENE-trending lateral ramp in the Bannock detachment fault (Fig. 15). The depth of the basin west of Rattlesnake Ridge decreases northward across this lateral ramp according to the gravity data and in accordance with our structural interpretation. The lateral ramp is also approximately marked by a change in the lithology of fault blocks (horses) along the detachment fault. North and south of the ramp, brecciated horses of Camelback Mountain Formation crop out, whereas within the ramp, horses of Blacksmith Formation crop out. These changes may be due to three-dimensional detachment fault geometries up-dip of the current exposures. The lateral ramp is folded by a W-plunging syncline-anticline pair that tightens the bends at the top and bottom of the ramp.

We are on the west limb of the Oxford Ridge anticline. The axial trace is at the narrows of Fivemile Canyon and is well exposed along Rattlesnake Ridge. This broad, open fold is subparallel to the Bannock detachment system and formed as an isostatic response to unloading above it. This broad anticline tilted portions of the detachment fault into unfavorable geometries for continued slip and likely initiated the process of excision. The transitional member of the Salt Lake Formation crops out near this stop to the N and NW for several hundred meters. The base and top of the transitional member can be seen to the south as the two high west-dipping ridges of zeolitized green to white tuffaceous siltstone beds (Fig. 7). There were several broad west-trending growth folds active during deposition of the transitional and lower Third Creek members south of here. Mileage Interval (Cumulative) Description 1.5

64.7

Continue westward along the Fivemile Creek road, and then climb up to a low ridge. Pass a Y-intersection at mile 64.2 and continue north for another 0.5 mile. Park and walk south to the top of the nearby small hill.

Stop 7: Faulted Hanging Wall above the Northern, Upper Flat in the Detachment Fault; UTM 0413284E 4663245N The goals of this stop are to (1) observe the northward truncation of hanging wall rocks against the lateral ramp; (2) discuss the style of hanging wall deformation and how it changes across the lateral ramp; (3) document a large en echelon step of the West Cache fault zone; and (4) observe some lithologies of the Third Creek Member. We are located on the upper flat of the Bannock detachment system where the Third Creek Member of the Salt Lake Formation is juxtaposed against Pocatello Formation. Looking south you can clearly see the two prominent zeolitized tuffaceous marker beds at the top and bottom of the transitional member striking north and being truncated by the detachment fault. Looking south again, we observe the thickest unfaulted section of Salt Lake Formation within the Malad and Bannock ranges. This thick unfaulted section above the southern, lower flat differs from the internally extended rocks above the northern upper flat at this stop. We suggest that increased faulting above the upper flat is due to the substantial northward thinning of the hanging wall across the Fivemile lateral ramp. To the west, the Clarkston-Junction strand of the West Cache fault lies along the eastern edge of the Dry Creek horst. Here it places the Third Creek Member of the Salt Lake Formation against Neoproterozoic to Lower Paleozoic rocks (Figs. 7 and 4). The West Cache fault is an active E-dipping Basin and Range normal fault that begins in southern Cache Valley, has an overall NNW strike, and a length of ~93 km (Solomon, 1999; Black et al., 2000). Our mapping shows that the Clarkston-Junction strand is the northern continuation of the West Cache fault (Steely and Janecke, 2005)

Evolution of a late Cenozoic supradetachment basin and the Steel Canyon fault is not part of the West Cache fault zone as Janecke and Evans (1999) had inferred. At this latitude, two other Basin and Range normal faults are also present: the W-dipping Deep Creek fault (DCF) on the west side of the Clifton horst and the E-dipping Dayton-Oxford fault (DOF) to the east. Both of these faults have along-strike overlap with the Clarkston-Junction strand of the West Cache fault (DCF ~15 km; DOF ~27 km). Gravity studies show a ~7 mGal gradient across the Clarkston-Junction fault in the northern part of the Weston Canyon quadrangle that increases southward to ~11 mGal (Fig. 15). Mapping shows that the Clarkston-Junction strand of the West Cache fault is linked to the Deep Creek fault by a network of normal faults that obliquely cross the Dry Creek horst. Most of the strain on the Clarkston-Junction strand of the West Cache fault steps en echelon eastward to the Dayton-Oxford fault. Thus the West Cache fault zone defines the western side of Cache Valley for its entire length. On the way back to the vehicles, detour west down the hill to examine outcrops of conglomerate beds and tufa heads within the Third Creek Member. The conglomerates at this location consist of pebbles to rare small cobbles dominated by green to white recycled tuffaceous siltstone and limestone clasts from the Salt Lake Formation and colored quartzites from the Brigham Group. Interbedded limestone, sandstone, and tephra are common in this interval. Limestone beds may be either tan silicified laminated to bedded micrite or gray with local tufa. Several gray tufa heads 10–50 cm in diameter crop out on the western slope of this hill. These lithologies suggest that within the lower Third Creek Member at this locale, deposition of coarse clastics occurred in a freshwater shallow lacustrine to near-shore setting. Mileage Interval (Cumulative) Description 0.4

65.1

3.0

68.1

0.2

68.3

Turn around and drive south to the Yintersection. Take the right fork down into the valley of Weston Creek. This road becomes 7200 W. and has an intersection with 300 South near some farm houses at mile 67—continue straight past the intersection. The road makes a left turn followed by a sharp right turn and another left. At the second right turn, a small dirt track joins the main road. Turn left (NE) onto this road. Continue 0.1 mile and park along the side of the road. Take care not to drive across the grass; this stop is on private property. Walk ~100 m north to a prominent W-dipping conglomerate bed.

Stop 8: Third Creek Member of the Salt Lake Formation and Evidence for Breakup; UTM 0413203E 4659050N At this stop we will examine Third Creek Member conglomerate beds, discuss evidence for an unroofing sequence in the

195

Tertiary basin-fill, and suggest an updated model of overall basin evolution. The conglomerate beds at this stop are typical of those observed throughout the Third Creek Member in the Weston Canyon and neighboring quadrangles. Although conglomerates account for less than ~20% of the entire Third Creek Member in the Weston Canyon quadrangle, they are more resistant than the under- and overlying calcareous to tuffaceous sandstone and siltstone, tephra, limestone, and localized tufa and form prominent ridges. Clast count data from the Clifton, Henderson Creek, Malad City East, and Weston Canyon quadrangles show that 52% of clasts are from the Brigham Group, 32% are from Paleozoic rocks, and 16% are recycled tuffaceous siltstone and limestone from the Salt Lake Formation (n = 68) (Fig. 9). The sudden influx of recycled Salt Lake clasts during the Third Creek Member suggests that some parts of the supradetachment basin were being uplifted and eroded while other areas experienced continued deposition. Recycled Salt Lake clasts, coupled with the change from widespread lacustrine deposition to more aerially-restricted subbasins record internal hanging wall breakup of the Bannock supradetachment basin (Fig. 11). The major phase of breakup began with deposition of the Third Creek Member, although growth folds active during the underlying transitional member suggest incipient breakup just prior to major deposition of conglomerates in the Weston Canyon quadrangle. Mileage Interval (Cumulative) Description 0.2

68.5

2.9

71.4

1.5

72.9

Retrace route back to the main dirt road and turn left (south). Drive 0.5 mi to the 4-way intersection (before the stop sign) and turn left (east). Continue on this gravel road, past a sharp right turn at mile 69.9. Road turns east. Make a hard left (north) onto a smaller dirt road between agricultural fields. Drive north 1.5 mi and park at the top of a small hill. This road may be impassable when wet.

Stop 9: Exposure of the Bannock Detachment Fault; UTM 0415888E 4658473N This stop requires a short walk to visit exposures of slickenlines on the Clifton strand of the Bannock detachment fault on top of a thin brecciated fault block. Follow the route from Figure 7 east to the wash and up the other side through the juniper trees. The coordinates of the exposure are UTM 0416118E 4658251N. The slickenlines are developed between the eroded lower Cache Valley Member of the Salt Lake Formation and a fault-bounded horse of brecciated Camelback Mountain Quartzite (Fig. 10C). Faulted below the Camelback Mountain Quartzite is the Scout Mountain Member of the Pocatello Formation. Up to 6 km of stratigraphic section is omitted along this fault.

A.N. Steely et al.

196

Previous workers suggested that the Bannock detachment fault slipped W or WSW based on fault and fold geometries in the hanging wall (Janecke and Evans, 1999; Carney and Janecke, 2005). The slickenlines preserved here and ~1.5 km south along this same fault, coupled with data from faults within horses along the detachment fault, show that the Bannock detachment fault system has a WSW-ENE extension direction, consistent with prior analyses of fault geometries. The UTM coordinates of the second slickenline exposure are 0416320E 4656903N. Return to the vehicles and turn around. Mileage Interval (Cumulative) Description 1.6

74.5

1.6

76.1

4.0

80.1

19.8

99.9

Retrace your route to the main gravel road. Turn left. Drive southeast to a T-intersection. Turn right (south) onto Franklin County D1. Proceed south through the town of Weston well into Utah, where the road becomes Utah Route 23. Continue on Route 23 for 19.8 mi. Turn right (east) onto Hwy 30 and drive 11.7 mi to the intersection with I-15.

SUMMARY Flat-on-flat relationships across ~6 km of the Bannock detachment fault show that the fault was active within ~1.6 to ~4.3 km of Earth’s surface yet had a negligible dip. A ~1-kmhigh lateral ramp in the detachment fault separates a highly faulted thin hanging wall above the upper flat from an intact and unfaulted hanging wall above the lower flat. Facies relationships in the synextensional Salt Lake Formation indicate a widespread lacustrine system from ca. 10.2 to ca. 9.6 Ma as the hanging wall of the detachment slipped as a single intact fault block during a translation phase. Lacustrine deposition continued but changed in character at ca. 9.6 Ma as the hanging wall began to break up into individual fault blocks and subbasins. West-trending growth folds, recycled Late Cenozoic basin fill, and renewed coarse clastic input record the reorganization of the basin during the breakup phase. Development of steeper Basin and Range–style normal faults postdate 4.4 Ma tephra beds within the Salt Lake Formation (Janecke et al., 2003). These faults dismembered and exposed the Bannock detachment system and form the modern topography in the area. ACKNOWLEDGMENTS Supported by four EDMAP (educational mapping component of the National Cooperative Geologic Mapping Program) grants, the Petroleum Fund of the American Chemical Society, an AAPG Grant-in-Aid, and the J.S. Williams fund of the Geology Department at Utah State University.

REFERENCES CITED Allmendinger, R.W., Sharp, J.W., Von Tish, D., Serpa, L., Brown, L., Kaufman, S., Oliver, J., and Smith, R.B., 1983, Cenozoic and Mesozoic structure of the eastern Basin and Range province, Utah, from COCORP seismic-reflection data: Geology, v. 11, p. 532–536, doi: 10.1130/00917613(1983)11<532:CAMSOT>2.0.CO;2. Axen, G.J., Fletcher, J.M., Cowgill, E., Murphy, M., Kapp, P., MacMillan, I., Ramos, V.E., and Aranda, G.J., 1999, Range-front fault scarps of the Sierra El Mayor, Baja California; formed above an active low-angle normal fault?: Geology, v. 27, no. 3, p. 247–250, doi: 10.1130/00917613(1999)027<0247:RFFSOT>2.3.CO;2. Biek, R.F., Oaks, R.Q., Jr., Janecke, S.U., Solomon, B.J., and Barry-Swenson, L.M., 2003, Geologic maps of the Clarkston and Portage quadrangles, Box Elder and Cache counties, Utah, and Franklin and Oneida counties, Idaho: Utah Geological Survey Map 194, 3 plates, scale 1:24,000, 41 p. Black, B.D., Giraud, R.E., and Mayes, B.H., 2000, Paleoseismic investigation of the Clarkston, Junction Hills, and Wellsville faults, West Cache fault zone, Cache County, Utah: Paleoseismology of Utah Volume 9: Utah Geological Survey Special Study 98, 23 p., 1 plate. Brady, R., Wernicke, B., and Fryxell, J., 2000, Kinematic evolution of a largeoffset continental normal fault system, South Virgin Mountains, Nevada: Geological Society of America Bulletin, v. 112, p. 1375–1397, doi: 10.1130/0016-7606(2000)112<1375:KEOALO>2.0.CO;2. Buck, W.R., 1988, Flexural rotation of normal faults: Tectonics, v. 7, p. 959–973. Carney, S.M., 2002, Evolution of a Miocene-Pliocene low-angle normal-fault system in the southern Bannock Range, southeast Idaho [M.S. thesis]: Logan, Utah State University, 177 p. Carney, S.M., and Janecke, S.U., 2005, Excision and the original low dip of the Miocene-Pliocene Bannock detachment system, SE Idaho: Northern cousin of the Sevier Desert detachment?: Geological Society of America Bulletin, v. 117, p. 334–353, doi: 10.1130/B25428.1. Carney, S.M., Janecke, S.U., Oriel, S.S., Evans, J.C., and Link, P.K., 2002, The Geologic map of the Clifton quadrangle, Franklin and Oneida Counties, Idaho: Idaho Geological Survey Technical Report T-03-4, scale 1:24,000, 2 plates, http://www.idahogeology.com/Products/reverselook.asp?switch =title&value=Geologic_Map_of_the_Clifton_Quadrangle,_Franklin_ and_Oneida_Counties,_Idaho. Chang, W., and Smith, R.B., 2002, Integrated seismic-hazard analysis of the Wasatch Front, Utah: Bulletin of the Seismological Society of America, v. 92, p. 1904–1922, doi: 10.1785/0120010181. Coney, P.J., and Harms, T.K., 1984, Cordilleran metamorphic core complexes; Cenozoic extensional relics of Mesozoic compression: Geology, v. 12, p. 550–554, doi: 10.1130/0091-7613(1984)12<550:CMCCCE>2.0.CO;2. Coogan, J.C., 1992, Structural evolution of piggyback basins in the WyomingIdaho-Utah thrust belt, in Link, P.K., Kuntz, M.A., and Platt, L.B., eds., Regional geology of eastern Idaho and western Wyoming: Geological Society of America Memoir 179, p. 55–82. Coogan, J.C., and DeCelles, P.G., 1996, Extensional collapse along the Sevier Desert reflection, northern Sevier Desert basin, western Unites States: Geology, v. 24, p. 933–936, doi: 10.1130/0091-7613(1996)024<0933: ECATSD>2.3.CO;2 Davis, G.A., and Lister, G.S., 1988, Detachment faulting in continental extension: perspectives from the southwestern U.S. Cordillera, in Clark, S.P., Jr., Burchfiel, B.C., and Suppe, J., eds., Processes in continental lithospheric deformation: Geological Society of America Special Paper 218, p. 133–159. DeCelles, P.G., 1994, Late Cretaceous-Paleocene synorogenic sedimentation and kinematic history of the Sevier thrust belt, northeast Utah and southwest Wyoming: Geological Society of America Bulletin, v. 106, p. 32–56, doi: 10.1130/0016-7606(1994)106<0032:LCPSSA>2.3.CO;2. Evans, J.P., and Oaks, R.Q., Jr., 1996, Three-dimensional variations in extensional fault shape and basin form: The Cache Valley basin, eastern Basin and Range province, United States: Geological Society of America Bulletin, v. 108, p. 1580–1593, doi: 10.1130/0016-7606(1996)108<1580: TDVIEF>2.3.CO;2. Evans, J.C., Janecke, S.U., and Oriel, S.S., 2000, Geologic map of the Malad City east quadrangle, southeast Idaho [unpublished M.S. thesis map]: Logan, Utah State University, scale 1:24,000. Eversaul, M.L., 2004, Basin structure of proposed Late Miocene to Pliocene supradetachment basins in SE Idaho, based on detailed gravity and geologic data [M.S. Thesis]: Pocatello, Idaho State University, 142 p.

Evolution of a late Cenozoic supradetachment basin Fanning, M.C., and Link, P.K., 2004, U-Pb SHRIMP ages of Neoproterozoic (Sturtian) glaciogenic Pocatello Formation, southeastern Idaho: Geology, v. 32, p. 881–884, doi: 10.1130/G20609.1. Fedo, C.M., and Miller, J.M.F., 1992, Evolution of a Miocene half-graben basin, Colorado River extensional corridor, southeastern California: AAPG Bulletin, v. 104, p. 481–493. Fowler, T.K., Jr., Friedmann, S.J., Davis, G.A., and Bishop, K.M., 1995, Twophase evolution of the Shadow Valley Basin, southeastern California: A possible record of footwall uplift during extensional detachment faulting: Basin Research, v. 7, p. 165–179. Friedmann, S.J., and Burbank, D.W., 1995, Rift basin and supradetachment basins: intracontinental extensional end-members: Basin Research, v. 7, p. 109–127. Gans, P.B., and Miller, E.L., 1983, Style of Mid-Tertiary extension in east-central Nevada, in Guidebook, Part 1, Geological Society of America Rocky Mountain and Cordilleran Sections Meeting: Utah Geological and Mining Survey Special Studies, v. 59, p. 107–160. Goessel, K.M., 1999, Tertiary stratigraphy and structural geology, Wellsville Mountains to Junction Hills, north-central Utah [M.S. thesis]: Logan, Utah State University, 231 p. Goessel, K.M., Oaks, R.Q., Jr., Perkins, M.E., and Janecke, S.U., 1999, Tertiary stratigraphy and structural geology, Wellsville Mountains to Junction Hills, north-central Utah, in Spangler, L.E., and Allen, C.J., eds., Geology of Northern Utah and Vicinity: Utah Geological Association Publication 27, p. 45–69. Henry, C.D., and Perkins, M.E., 2001, Sierra Nevada—Basin and Range transition near Reno, Nevada: Two stage development at ca. 12 and 3 Ma: Geology, v. 29, p. 719–722, doi: 10.1130/0091-7613(2001)029<0719: SNBART>2.0.CO;2. Janecke, S.U., 2004, Translation and breakup of supradetachment basins: Lessons from the Grasshopper, Horse Prairie, Medicine Lodge, Muddy Creek, and Nicholia Creek basins SW Montana (abs.): Geological Society of America Abstracts with Programs, v. 37, no. 5, p. 546. Janecke, S.U., and Blankenau, J.C., 2003, Extensional folds associated with Paleogene detachment faults in SE part of the Salmon basin: Northwest Geology, v. 32, p. 51–73. Janecke, S.U., and Evans, J.C., 1999, Folded and faulted Salt Lake Formation above the Miocene to Pliocene(?) New Canyon and Clifton detachment faults, Malad and Bannock ranges, Idaho: Field trip guide to the Deep Creek half graben and environs, in Hughes, S.S., and Thackray, G.D., eds., Guidebook to the Geology of Eastern Idaho: Pocatello, Idaho, Idaho Museum of Natural History, p. 71–96. Janecke, S.U., VanDenburg, C.J., Blankenau, J.C., and M’Gonigle, J.W., 2000, Long-distance longitudinal transport of gravel across the Cordilleran thrust-belt of Montana and Idaho: Geology, v. 28, p. 439–442. Janecke, S.U., Carney, S.M., Perkins, M.E., Evans, J.C., Link, P.K., Oaks, R.Q., Jr., and Nash, B.P., 2003, Late Miocene-Pliocene detachment faulting and Pliocene-Pleistocene Basin-and-Range extension inferred from dismembered rift basins of the Salt Lake Formation, southeast Idaho, in Raynolds, R., and Flores, R., eds., Cenozoic Paleogeography of the Rocky Mountains: Society for Sedimentary Geology (SEPM) Special Publication, p. 369–406. John, B.E., 1987, Geometry and evolution of a mid-crustal extensional fault system: Chemehuevi Mountains, southeastern California, in Coward, M.P., Dewey, J.F., and Hancock, P.L., eds., Continental extensional tectonics: London, Geological Society Special Publication 28, p. 313–335. Kruger, J.M., Crane, T.J., Pope, A.D., Perkins, M.E., and Link, P.K., 2003, Structural and stratigraphic development of Neogene basins in the Lava Hot Springs and March Valley areas, southeast Idaho: Two phases of extension, in Raynolds, R., and Flores, R., eds., Cenozoic Paleogeography of the Rocky Mountains: Society for Sedimentary Geology (SEPM) Special Publication, p. 407–457. Link, P.K., 1982a, Geology of the upper Proterozoic Pocatello Formation, Bannock Range, southeastern Idaho [Ph.D. thesis.]: Santa Barbara, University of California, 131 p. Link, P.K., 1982b, Structural Geology of the Oxford Peak and Malad Summit quadrangles, Bannock Range, southeastern Idaho, in Powers, R.B., ed., Geologic studies of the Cordilleran thrust belt: Denver, Colorado, Rocky Mountain Association of Geologists, p. 851–858. Link, P.K., Jansen, S.T., Halimdihardja, P., Lande, A.C., and Zahn, P.D., 1987, Stratigraphy of the Brigham Group (Late Proterozoic-Cambrian), Bannock, Portneuf, and Bear River Ranges, southeastern Idaho, in Miller,

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Printed in the USA

Geological Society of America Field Guide 6 2005

Utah’s state rock and the Emery coalfield: Geology, mining history, and natural burning coal beds Glenn B. Stracher Division of Science and Mathematics, East Georgia College, Swainsboro, Georgia 30401, USA David E. Tabet Utah Geological Survey, P.O. Box 146100, 1594 West North Temple, Salt Lake City, Utah 84114, USA Paul B. Anderson Consulting Geologist, 807 East South Temple, #101, Salt Lake City, Utah, 84102, USA J. Dénis N. Pone School of Geosciences, University of Witwatersrand, Johannesburg, P. Bag X 3, Wits 2050, South Africa

ABSTRACT The commercial development of Utah’s state rock, coal, by the Union Pacific and Denver and Rio Grande Western railroads began in the late 1800s. Ninety-five percent of the coal produced in Utah today is used to generate electricity, and recent developments in coal-bed methane characterization and extraction are promising for the state’s economy. The Emery coalfield in central Utah consists of sediments deposited along the western margin of the Cretaceous Western Interior Seaway and subsequently deformed during the Laramide orogeny and Basin and Range deformation. The coalfield contains important methane and bituminous deposits in the Ferron Sandstone Member of the Cretaceous Mancos Shale. Ferron coal beds exposed on the northwest-dipping western flank of the San Rafael Swell have been burning for decades. The collapse of strata overlying the Ferron as a consequence of burning provides conduits for the circulation of oxygen, thereby promoting the smoldering that can be observed today. Dating clinker in the Emery coalfield would provide useful information about the timing of certain geologic events in Utah on a local and perhaps regional scale. Federal agencies and laws help safeguard coal mining in the United States today. The environmental effects of Utah’s natural burning and mine-related coal fires are unknown. However, such fires elsewhere are responsible for pollution, including acid rain, and they are responsible for a variety of human diseases. Keywords: Emery coalfield, Ferron Sandstone Member, Mancos Shale, coal fires, natural burning coal fires.

Stracher, G.B., Tabet, D.E., Anderson, P.B., and Pone, J.D.N., 2005, Utah’s state rock and the Emery coalfield: Geology, mining history, and natural burning coal beds, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 199–210, doi: 10.1130/ 2005.fld006(09). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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The Emery coalfield in central Utah (Fig. 1) trends southwest-northeast across four counties: Sevier, Sanpete, Emery, and Carbon. Bituminous-coal beds in the Upper Cretaceous Ferron Sandstone Member of the Mancos Shale are noteworthy for their methane and hard coal production (Montgomery et al., 2001). The Ferron is exposed on the northwest-dipping western flank of the San Rafael Swell. It consists of littoral-shelf sandstones in the east overlain by transgressive-regressive fluvial-deltaic sediments that progressively thicken to the west, extending beneath the Wasatch Plateau. The Emery mine, the most productive in the Emery coalfield since 1970 (Table 1), is the only active mine in the coalfield today. It is in the bituminous-I coal bed of the Ferron. Part of Utah’s coal legacy includes evidence of extensive natural burning coal fires that resulted from spontaneous combustion and lightening strikes, as well as some due to mining activities; some of these fires are still active today.

This paper presents a synopsis of the geology of the Ferron, a historical overview of mining in the Emery coalfield, and illustrations of baked coal beds in the Ferron at the southern end of this coalfield. In addition, the field trip agenda includes demonstration of the instrumentation and hands-on experience in collecting coal-fire data in the field. GEOLOGY Stratigraphy The Mancos Shale in Utah is a sequence of Cretaceous marine shales intercalated with fluvial-deltaic sandstones and assorted continental-sedimentary rocks deposited along the western margin of the Cretaceous Western Interior Seaway (Fig. 2). The seaway extended across the interior of North America from Mexico to the Arctic and covered eastern and central Utah. The fluvial-deltaic sandstones consist of stream-transported sediment

Figure 1. Maps of the field trip route following I-15 south of Salt Lake City, Utah, to the six stops (S1–S6) in the Emery coalfield.

Utah’s state rock and the Emery coalfield eroded from Paleozoic and basement rocks of the CretaceousTertiary Sevier orogenic belt, west of the seaway (Ryer, 1981, 1991; Gardner and Cross, 1994; Montgomery et al., 2001). The foreland depositional basin containing the Cretaceous sediments is asymmetrical in cross section because it subsided faster along its western margin in response to depositional loading associated with erosion of the Sevier belt. Sandstones, conglomerates, and gravels of the Indianola Group, eroded by streams from the Sevier orogenic belt, were deposited west of the foreland basin, and are time-stratigraphic equivalent with the Mancos Shale. The Ferron Sandstone Member of the Mancos Shale is the focus of this trip; it is Turonian-Coniacian in age (Schwans and Campion, 1997; Fig. 2). The Clawson and overlying Washboard strata that comprise the lowermost section of the Ferron are finegrained littoral-shelf sandstones up to 100 ft (~30 m) thick, only found along the eastern end of the Ferron. The upper Ferron consists of transgressive-regressive fluvial-deltaic sandstones, shales, siltstones, and coal beds varying in total thickness from 180 ft (55 m) near Emery up to 755 ft (230 m) beneath the Wasatch Plateau. The Ferron contains 12 coal zones starting with “A” at the bottom and proceeding stratigraphically up to “L” at the top, but only 8 are developed (Lupton, 1916). Up to 13 coal beds may occur within a zone, and individual beds vary in thickness from 1 to 10 ft (~0.3–3 m). Most coal beds contain one or more tonsteins. The coal beds pinch-out to the east into littoral and deltaic-front sandstones that are overlain by marine shales and siltstones (Ryer 1981; Gardner et al., 1992). Westward, the coal beds pinch-out into fluvial, floodplain, and upper deltaic sandstones and shales (Ryer, 1981). Structure The Ferron is exposed for ~35 mi (56 km) along strike on the shallow-dipping western flank of the San Rafael Swell, a southwest-northeast–trending anticline formed during the Laramide (Late Cretaceous–Eocene) orogeny. West of the outcrop area, the

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TABLE 1. EMERY COALFIELD MINES Name of coal mine Emery Browning Dog Valley Sun Valley Cowboy Peterson Moore Williams Cox Bear Gulch Casper Willow Springs Willow Basin

Years of operation

Coal seam mined

1970–2005 1881–1970 1930–1990 1970–1973 1900–1920 1935–1938 ca. 1905 Before 1916 Before 1916 1897–1916 Before 1916 1932–1946 Before 1910

I I I I I I or J I or J I or J I or J C C A A

Cumulative production* 7,651,344 1,318,000 649,000 76,800 1000 4000 1500 700 15 ? 1500 16,000+ ?

Note: Modified from Quick et al., 2004. *Through 2004.

Ferron dips beneath the surface, and according to bore hole data, its coal beds thin and pinch-out under the Wasatch Plateau (Tabet et al., 1995; Montgomery et al., 2001). On the western flank of the Swell, the Ferron is locally folded and reverse faulted—structural features thought to be genetically related to the San Rafael uplift (Tripp, 1989; Burns and Lamarre, 1997). Farther west, borehole data reveal that during the Miocene to Recent, the Ferron was normal faulted (Fig. 2) by Basin and Range deformation (Doelling, 1972; Montgomery et al., 2001). To the north, the Ferron plunges beneath the Book Cliffs, the southern edge of the Uinta Basin. To the south, it extends beneath Oligocene latite lavas of the Fish Lake Plateau that range in thickness from 400 to 1000 ft (122–305 m) (Hintze, 1988). MINING HISTORY The economic importance of coal to Utah for over 150 years was acknowledged by an act of the Utah State Legislature on

Figure 2. Schematic cross section illustrating strata in the western margin of the Cretaceous Western Interior Seaway that covered eastern and central Utah. Indianola Group sediments are timestratigraphic equivalent with the Mancos Shale. Normal faults and graben development associated with Basin and Range deformation and reverse fault (through the Ferron) associated with the San Rafael Swell are shown without offset. Modified from Armstrong (1968) and Montgomery et al. (2004).

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March 13, 1991, officially recognizing coal as the Utah state rock (Thatcher, 1994). According to 2003–2004 statistics of the National Mining Association (2005a), Utah is the thirteenth largest coal producing state in the United States. In addition, since the 1990s, technological advances in coal-bed methane characterization and extraction have made the northern Emery coalfield in Emery and Carbon counties one of the most profitable gas fields in North America, with wells averaging >500 million ft3/ day (Montgomery et al., 2001). Coal was mined in Utah by white settlers as early as the 1850s and used as a heating and cooking fuel in lieu of wood, which was scarce and used primarily for building purposes (Taniguchi, 1990). However, commercial coal mining in the state did not begin until around 1870. Today, as in the past, coal mining in Utah is predominantly conducted by underground methods, whereas in the rest of the United States, nearly two-thirds of coal production is by surface mining (National Mining Association, 2005b). In the late 1800s, the Union Pacific and Denver and Rio Grande Western railroads were instrumental in the commercial development of coal mining in Utah and settlement in the sparsely populated western United States. The scarcity of western settlers created a labor shortage, so the railroads hired agents who recruited foreign immigrants from China, Italy, Mexico, and elsewhere to work in Utah’s coal mines. Conditions in the early coal mines were often poor, as mining methods were still primitive, and included working with open-flame lamps, hand cutting and loading of the coal, and uncontrollable coal dust–related explosions. In addition, many immigrant miners were initially unskilled, experienced communication difficulties amongst each other as well as with company officials, and they lived in railroad-owned company towns where prices were high for the necessities of life. Throughout the second half of the 1800s and early 1900s, numerous unsuccessful mining strikes occurred in response to poor working and objectionable living conditions. Miners wanted the right to representation by organized labor unions, especially the United Mine Workers of America (UMWA), which was active in the eastern states (Taniguchi, 1990; Powell 1994). In spite of these strikes and concomitant violent outbursts, conditions generally did not improve until the U.S. Congress passed the National Labor Relations Act (Wagner Act) of 1935. Part of Franklin D. Roosevelt’s New Deal policy, this act established the right of labor to collective bargaining in organized unions with the companies (Roosevelt, 1938, p. 294–295). Coal mining in the United States today is dramatically different from early mining, employing highly efficient machinery to cut and load the coal to relieve workers of back-breaking work, using water sprays to control harmful coal dust and diluting it with powdered limestone to make it non-explosive, and providing safe roof support and ventilation. The Office of Surface Mining (OSM) regulates surface disturbances of coal mining, and the Mine Safety and Health Administration (MSHA) regularly inspects coal mines to ensure the use of safe practices. Federal laws that affect the coal-mining industry today include the following:

• The National Historic Preservation Act (1966), which governs the preservation of historic properties throughout the United States; • The National Environmental Policy Act (1969), which established a national policy for the environment; • The Endangered Species Act (1973), which governs the protection of endangered species; • The Resource Conservation and Recovery Act (1976), which governs the control of hazardous waste; • The Clean Water Act (1977), which regulates the discharge of pollutants into water; and • The Clean Air Act (1990), which regulates the discharge of pollutants into the air. The coal-mining industry in Utah remains an important contributor to the economic health of the state. Almost three billion short tons of recoverable reserves in Utah account for ~1.03% of the U.S. reserve base (OSM, 1997; U.S. Department of Energy, 2004). Ninety-five percent of Utah’s 24.6 million short tons of annual coal production from 331 million short tons of active mine reserves is used to generate electricity. In addition, millions of dollars worth of coal is exported overseas by Utah every year (Tabet et al., 1998; OSM, 2004; U.S. Department of Energy, 2004). The Emery Coalfield Coal was first mined in the southern Emery coalfield in 1881. At least eight mines serving local consumers were opened during the next 50 years. During the 1930s, the Willow Springs, Peterson, Dog Valley, and Browning mines led the way in hard coal production, which remained nearly continuous until 1990 (Quick et al., 2004). The only active mine in the coalfield today is the Consolidation Coal Company’s Emery mine, near the older Browning Mine, operating since 1970 in the bituminous-I coal seam. Emery has been the most productive mine in the coalfield (Table 1) even though I-seam production over the years has varied in response to fluctuating market prices, contracts, and available rail transport. As of 2004, the cumulative Emery coalfield production was slightly over 9.8 million short tons from an original recoverable reserve of 429 million (Brill et al., 2004). COAL FIRES IN UTAH In the early days of mining in Utah, coal-mine explosions and fires triggered by a variety of mine-related activities claimed the lives of numerous people in the state’s coalfields. Although coal-mine fires today are an ever-present danger, they are less common. A National Institute for Occupational Safety and Health study of coal-mine fires from 1990 through 1999 listed only one underground coal-mine fire in Utah for that period (De Rosa, 2004). The most recent coal-mine fire occurred in July 2000 at the now-closed Willow Creek mine in the Book Cliffs coal field to the north of Price, Utah. The fire in this gassy mine apparently started when falling rock caused a spark that ignited a pocket of built-up methane gas near the longwall mining machine, and the

Utah’s state rock and the Emery coalfield fire then spread explosively throughout the mine (Nichols, 2001). The mine was flooded and the fire extinguished, but unfortunately two lives were lost as a result of the incident. The mine has been permanently closed since the fire. Within the last several years, more than a dozen coal fires were reported burning in Utah, several as a consequence of mining but most from natural causes (Bauman, 2002). Naturally occurring fires are locally burning in a number of Utah’s 22 coalfields. These include coal beds in the Emery coalfield, where 13 beds in the Ferron are exposed in the southern part of the field due to uplift and erosion of the Mancos Shale. These exposures occur along the Emery-Sevier County line and along the field trip route in southwestern Emery County. Extensive natural burning has occurred most frequently along southern slopes and is characterized by small, isolated areas of smoldering coal beds and extensive baked zones in the adjacent country rock. The baked rocks (clinker) vary in color from red to yellow-orange, depending on their distance from and the intensity of burning in the coal beds. The colors are most intense adjacent to the burning seams and grade into the tan, gray, and black unaltered rocks interbedded with the coal. Previous and currently active natural burning coal fires in the Emery coalfield are likely the result of either spontaneous combustion due to oxidation and self heating within the coal or lightening strikes. When northwest-dipping Ferron coals were exposed at the surface on the western flank of the San Rafael Swell due to erosion of the overlying Cretaceous rocks, spontaneous combustion or lightening strikes ignited the beds. As burning occurred, the collapse of overlying strata provided conduits for the circulation of oxygen and hence continued burning. Today, there are isolated areas of smoldering on the surface. FIELD TRIP STOPS IN THE EMERY COALFIELD The directions to and global positioning satellite locations (latitude, longitude, and elevation) of the six field trip stops (Fig. 1) in the Emery coalfield in Emery County, Utah, are given below, along with a description and illustration of each stop. In the field trip descriptions and illustrations, “Ferron sandstone” and “Ferron No. 1, 2, 4, and 7 sandstone” refer to fluvial shoreline and marine sand, respectively. The marine sands are numbered from stratigraphically lowest to highest according to increasing integral value. The trip departs from the Salt Palace Convention Center in Salt Lake City. Stop 1 (Fig. 3): Ferron Sandstone Depositional History and Deformation Stop 1 is at 38°48′1′′N, 111°15′58.5′′W, 6100 ft (1859 m). From Salt Lake City, drive 121.1 mi (~195 km) south, following I-15 to the town of Scipio, and take Exit 188 onto U.S. Hwy 50. Follow Hwy 50 south for 28.3 mi (45.5 km) to Salina and then drive 2.5 mi (4.0 km), following the signs for the onramp to I-70. Next, follow I-70 east for 39.6 mi (63.7 km), before the intersec-

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tion of I-70 with State Highway 10, to the Emery coalfield. Park on the shoulder of the road. The Ferron sandstone along I-70 here records the history of deposition in a foreland basin. The complex stratigraphy is characterized by early filling of an asymmetrical basin and continued seaward progradation of the shoreline, then a period of vertical stacking of shorelines, and finally, back-stepping of the shorelines as the sea transgressed ~30 mi (~48 km) to the west. This stop is at a relatively landward position in the progradation of the Ferron and just short of where the transition to vertical aggradation begins. Looking north across I-70, several interesting features about the shoreface or parasequence units and overlying coal beds are observable (Fig. 3A). First, note the landward pinch-out of the C shoreface. This pinch-out also occurs at Ivie Creek, a short distance to the north, and in the cliffs a few miles to the south. The orientation of the pinch-out is north-south, atypical from the more common northwest-southeast shoreline orientation. Below the C shoreface are two vertically stacked wave-modified shorefaces, A and B. Above the three shorefaces and within the road cut are the lower delta plain facies with interbedded coal, shale, and lenticular channel sands. The first coal in the sequence is the A coal bed. Above the road cut and along the top of the photograph are alluvial-plain facies, representing more landward facies that prograded over the older and more seaward facies. The stratigraphically higher G coal bed is burned, with the characteristic red “bloom.” Although not shown in Figure 3A, by looking up the road cut and to the east, the C coal bed is clearly visible with associated lenticular channel sandstone. Disharmonic and parasitic interference folds formed during soft-sediment deformation of the Ferron sandstone occur at the top of the road cut along the south side of I-70 (Figs. 3B and 3C), adjacent to the shoulder of the road. The deformation is presumably the result of water-saturated sand slumping along a shoreface, likely related to seismic activity shortly after deposition. In addition, Oligocene volcanic boulders (25 Ma) (Hintze, 1988) from the Fish Lake Plateau unconformably overlie the Ferron here (Fig. 3D). As discussed earlier and illustrated in Figure 3A, three parasequence deltaic cycles of the Cretaceous Ferron-2 (Kf-2) parasequence set are exposed lower in the road cut, just down the hill from the disharmonic folding and volcanics. About half way through the road cut is the sub-A coal, which marks the boundary between the underlying Kf-1 parasequence set and the Kf-2 above. Stop 2 (Fig. 4): Emery Coal-Mine Portal Stop 2 is at 38°52′28′′N, 111°13′58′′W, 6087 ft (1855 m). From Stop 1, follow I-70 east for 3.4 mi (~5.5 km) and take the Ranch Exit. At the bottom of the exit ramp, turn left and drive along Miller Canyon Road (paved) for 7 mi (~11 km) and then turn left onto a dirt road, before the town of Emery. Drive 0.1 mi (~0.2 km), and turn left at the stop sign onto another dirt road. Drive 1.4 mi (~2.3 km) and bear to the left at the fork in the dirt

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Figure 3. (Stop 1). (A) Complex stratigraphy of the Ferron, including shorefaces (deltaic parasequences), discussed in the text. Modified from Anderson et al. (2003). (B) and (C) Disharmonic and parasitic-interference folds in the Ferron sandstone, formed during soft sediment deformation. Fold termini are outlined in solid black, hinge surfaces and interference folds in black dashed and dotted lines, respectively. (D) D.E. Tabet (left) and P.B. Anderson (right) standing on and looking upslope across the Ferron sandstone as J.D.N. Pone points to overlying Oligocene volcanic rock eroded and transported in from the Fish Lake Plateau, south of the Emery coalfield.

Utah’s state rock and the Emery coalfield road, continuing for 0.8 mi (~1.3 km) and across a paved road to the Emery mine portal. The Consolidation Coal Company’s Emery mine portal was developed in 2003, 2 mi (~3 km) northeast of the original portal. The new portal was necessary because access to unmined coal from the original portal would have required extensive rehabilitation of the old mine tunnels, which were unused for ~10 yr after the mine closed. The new portal was developed by digging an inclined tunnel down 50 ft (~15 m) to access the I coal bed. This bed is 10–15 ft (~3–5 m) thick in this area. Miners are only removing the lower 8 ft (~2.4 m) of the bed because the upper part is too high in sulfur to be marketed without coal cleaning. The coal mine conveyor illustrated in Figure 4 transports coal to the surface where it is crushed and sized, and the dump conveyor then stockpiles it. Next, it is transported by the surge bin conveyor into the surge or weigh bin, from which it is then loaded into trucks for transport. Stop 3 (Fig. 5): Ferron Sandstone Member and Overlook Stop 3 is at 38°51′08′′N, 111°12′58′′W, 6361 ft (1939 m). Continue past the Emery mine in Stop 2 for 1.8 mi (~2.9 km) and bear to the left at the fork in the dirt road. Drive for 0.3 mi (~0.5 km) and bear to the left at the next fork (the road to the right has a gate across it). Drive 0.2 mi (~0.3 km) along this road and park. The Ferron No. 4 sandstone caps the mesa-cliff top, as shown in Figure 5A. Looking up dip, and down section, the Dakota Sandstone (Fig. 2) and San Rafael Swell can be seen in the distance. Looking to the left, a mesa capped by Ferron and

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underlain by the less weather resistant Tununk Member of the Mancos Shale can be seen. Climb down 50 ft (~15 m) below the top of the cliff in the foreground of Figure 5A and walk 50 ft (~15 m) to the right to see large Ferron No. 4 sandstone boulders, which cover the former entrance to a coal mine. Continue 15 ft (~4.6 m) into the entrance of a tunnel: the A and C coal beds in addition to bedding planes and tafoni (cavernous weathering) in the Ferron No. 4 sandstone are visible (Figs. 5B and 5C). Stop 4 (Fig. 6): Gas Vent in the Ferron Sandstone Stop 4 is at 38°50′19′′N, 111°14′18′′W, 6423 ft (1958 m). From Stop 3, drive 0.2 mi (~0.3 km) back to the previous road fork and then 0.3 mi (~0.5 km) back down the road. Take a left turn at the fork and drive 1.1 mi (~1.8 km) past the red clinker on the right to another fork in the dirt road. Take a hard right turn here, drive 0.1 mi (~0.2 km), and then park. Walk ~138 ft (42 m) N80°E from the parking location. A Pasco thermocouple probe at a coal-fire gas vent located at 38°50′18′′N, 111°14′20′′W, 6416 ft (1956 m), in the Ferron sandstone will be used to demonstrate temperature data collection methods (Figs. 6A and 6B) to field trip participants. Continue ~488 ft (149 m) N48°E to the edge of Quitchupah Canyon located at 38°50′14′′N, 111°14′24′′W, 6380 ft (1945 m). A mesa can be seen across the canyon at S15°E (Fig. 6C). Extensive baking of the Ferron sandstone in the mesa and 20–30 ft (6–9 m) of subsidence of its top due to a volume reduction from burned out coal beds are readily observable. Stop 5 (Fig. 7): Reverse Fault and Sedimentary Structures Stop 5 is at 38°50′29′′N, 111°13′27′′W, 6477 ft (1974 m). Drive 0.1 mi (~0.2 km) back to the previous road fork described in the directions to Stop 4 and continue down the road for 1.03 mi (~1.66 km), driving past another road fork to the first intersection after this fork. Turn left and park. At this location, both unbaked and baked Ferron sandstone as well as baked nonmarine shale within the Ferron occur above and below the exposed I coal bed (Fig. 7A). Also visible here are areas where the baked Ferron sandstone has collapsed into voids left by completely burned I coal. The bedding plane offset between baked and unbaked Ferron sandstone reveals a reverse fault. This fault postdates the undetermined age of the Ferron clinker. In addition, the inclined Ferron sandstone layers above the road are possibly a lateral point bar fluvial deposit, collapsed into Ferron clinker (Fig. 7B). A laterally discontinuous channel deposit of lenticular Ferron sandstone is visible above the road as well. Stop 6 (Fig. 8): Coal-Fire Data Collection

Figure 4. (Stop 2). Consolidation Coal Company’s Emery mine portal, opened in 2003. The bituminous-I coal seam is ~50 ft (~15 m) down. The covered conveyor belts transport coal from the mine to the surface and into the surge bin for loading into trucks (see text). The dip slope of the Ferron sandstone is in the foreground, including the dirt road (light color), and the Wasatch Plateau is in the background.

Stop 6 is at 38°53′24′′N, 111°12′13′′W, 6255 ft (1907 m). Return to the intersection just before Stop 5, turn left, and drive 5.4 mi (~8.7 km) down the road, past a fork in the dirt road and the Emery mine, to Miller Canyon Road. Turn right here and drive

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Figure 5. (Stop 3). (A) Ferron No. 4 sandstone (looking east) caps the cliff top as shown. The Dakota Sandstone and San Rafael Swell appear up dip and down section in the distance. (B) Ferron No. 4 sandstone and coal beds exposed in a tunnel ~15 m (50 ft) below the top of the cliff in the foreground of (A). A shale parting occurs between the A and C coal beds. Ferron boulders and rubble cover the contact between the shale and the C coal bed. The A bed is underlain by carbonaceous shale. (C) Tafoni (cavernous weathering) and bedding plane surfaces exposed in the tunnel at (B).

1.2 mi (~1.9 km), then turn left on a dirt road, cross the cattle guard, and park. Hike N43°E across the Ferron No. 7 sandstone for 3 mi (~5 km) to the burn area at the edge of Muddy Canyon. The Ferron No. 7 can be seen here along with the actively burning J coal bed underneath (Figs. 8A and 8B). Field trip participants will be given the opportunity to collect temperature data and gas samples here for analyses. In addition, collection procedures for solid combustion by-products will be demonstrated using a metal spoon, spatula, and collection bottles. During this process, it is important to minimize contamination of the solid with the substrate it formed on, or else analyses and scanning electron microscope imaging of the desired material may be exceedingly difficult or impossible (Stracher, 2003). DISCUSSION Clinker in the Emery coalfield records the occurrence of natural burning coal fires that occurred in the geologic past. The

timing of such fires remains to be determined and likely occurred subsequent to the exposure of coal beds at the surface due to uplift and erosion. Fission-track, uranium-thorium/helium, and paleomagnetic dates of detrital zircons in clinker reveal that natural burning coal fires in the Powder River basin in Wyoming and Montana occurred during the past 4 m.y. (Heffern and Coates, 2004). Dating clinker in the Emery coalfield by analogous techniques would provide useful information about the timing of geologic events in Utah on a local and perhaps regional scale. For example, clinker offset along fault planes and exposed in canyon walls could be used to establish the maximum age of faulting and initial canyon development, respectively. The environmental effects of Utah’s natural burning and mine-related coal fires are unknown. However, some trace elements associated with, and mobilized during, combustion are known to promote soil, water, and air pollution and contribute to the formation of acid rain (Stracher, 2002, 2003; Stracher and Taylor, 2004). Such pollutants, including mercury, sele-

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Figure 6. (Stop 4). (A) G.B. Stracher using a Pasco Physics Explorer data logger and high temperature type K thermocouple probe at a South African mine fire. (B) G.B. Stracher (left), J.L. Stracher, and J.D.N. Pone (right) using the same kind of apparatus to measure the temperature (76 °C) at a more quiescent coal-fire gas vent in the Ferron sandstone. The Wasatch Plateau (looking west) is visible in the background. (C) Extensive baking of the Ferron in a mesa, looking S15°E across Quitchupah Canyon. An outcrop of partially burned I coal is exposed along the side of the mesa and the Tununk Member of the Mancos Shale is exposed at its base. Subsidence occurred in response to the volume reduction from burned out coal beds in the mesa.

nium, thallium, arsenic, fluorine, and organic compounds, have destroyed floral and faunal habitats, forced entire communities to relocate, and are responsible for an array of human diseases including thallium poisoning, hyperkeratosis (arsenic poisoning), dental and skeletal fluorosis (osteosclerosis), pulmonary heart disease, and lung cancer (Johnson et al., 1997, p. 19; World Resources Institute, 1999, p. 63–67; Finkelman et al., 1999, 2001, 2002; Finkelman, 2004). Analyses of coal-fire gas from Utah and trace element analyses of minerals formed as coal combustion by-products may reveal potential vectors for the transmission of toxins to humans by food grown in soils that contain these minerals or even by dust particles that are inhaled. Coal fires, like those observed in the Emery coalfield, destroy an important natural resource, and if left to burn, add greenhouse gases and a variety of toxins to the atmosphere. Enhanced understanding of spontaneous combustion processes and advances in firefighting techniques for extinguishing natural burning coal fires would help preserve valuable coal resources for beneficial use and reduce the flux of toxins and greenhouse gases into Earth’s global subsystems such as the atmosphere, biosphere, and lithosphere.

ACKNOWLEDGMENTS The authors thank the Coal Geology Division of the Geological Society of America for sponsoring this GSA field trip. Virginia H. Lewick of the Franklin D. Roosevelt Presidential Library and Museum in Hyde Park, New York, as well as Gina Strack of the Utah History Research Center and Paul E. Burrows of the Office of Information Technology at the University of Utah, Salt Lake City, assisted us with reference materials. We are grateful to the Utah Department of Natural Resources for providing a vehicle with which to prepare for this field trip. In addition, we thank Michael Hylland and Kimm Harty of the Utah Geological Survey and Joel L. Pederson of Utah State University for their review of this manuscript. REFERENCES CITED Anderson, P.B., Tabet, D.E., and Hampton, G.L., III, 2003, Coalbed gas deposits of central Utah: Rocky Mountain Section, American Association of Petroleum Geologists, Field Trip 18. Armstrong, R.L., 1968, Sevier orogenic belt in Nevada and Utah: Geological Society of America Bulletin, v. 79, p. 429–458. Bauman, J., 2002, CEU scientist wields tool to snuff coal fire dragons: Salt Lake City, Utah, Deseret News Publishing Company, http://deseretnews.com/ dn/view/0,1249,450023727,00.html (accessed 26 Aug. 2005).

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Figure 7. (Stop 5). (A) Baked Ferron sandstone and nonmarine shale within the Ferron Sandstone Member. Note the reverse fault in the Ferron sandstone and its collapse into the area where the I coal bed completely burned (also at knee level next to G.B. Stracher). (B) Baked area that includes (A). A laterally discontinuous channel of lenticular Ferron sandstone and possible point bar deposit are also visible.

Brill, T., Allred, J., and Vanden Berg, M.D., 2004, Annual review and forecast of Utah coal–—2003: Utah Energy Office of the Department of Natural Resources, 33 p. Burns, T.D., and Lamarre, R.A., 1997, Drunkards Wash Project: coalbed methane production from Ferron coals in east-central Utah: Tuscaloosa, University of Alabama, Proceedings of the International Coalbed Methane Symposium, Paper 9709, p. 507–520. De Rosa, M.I., 2004, Analysis of mine fires for all U.S. underground and surface coal mining categories—1990–1999: U.S. Department of Health and Human Services, National Institute for Occupational Safety and Health, Information Circular 9470, 36 p. Doelling, H.H., 1972, Central Utah coal fields: Sevier-Sanpete, Wasatch Plateau, Book Cliffs, and Emery: Salt Lake City, Utah Geological and Mineralogical Survey, Monograph Series, v. 3, 572 p. Finkelman, R.B., 2004, Potential health impacts of burning coal beds and waste banks, in Stracher, G.B., ed., Coal fires burning around the world: a global catastrophe: International Journal of Coal Geology, v. 59, no. 1–2, p. 19–24. Finkelman, R.B., Belkin, H.E., and Zheng, B., 1999, Health impacts of domestic coal use in China: Proceedings of the National Academy of Sciences of the United States of America, v. 96, p. 3427–3431, doi: 10.1073/ pnas.96.7.3427. Finkelman, R.B., Skinner, H.C., Plumlee, G.S., and Bunnell, J.E., 2001, Medical Geology: Geotimes, v. 46, no. 11, p. 21–23. Finkelman, R.B., Orem, W., Castranova, V., Tatu, C.A., Belkin, H.E., Zheng, B., Lerch, H.E., Marharaj, S.V., and Bates, A.L., 2002, Health impacts of coal and coal use: possible solutions: International Journal of Coal Geology, v. 50, p. 425–443, doi: 10.1016/S0166-5162(02)00125-8. Gardner, M.H., and Cross, T.A., 1994, Middle Cretaceous paleogeography of Utah, in Caputo, M.V., Peterson, J.A., and Franczyk, K.J., eds., Mesozoic systems of the Rocky Mountain region, USA: Rocky Mountain Section, Society for Sedimentary Geology (SEPM), p. 471–503.

Gardner, M.H., Barton, M.D., Tyler, N., and Fisher, R.S., 1992, Architecture and permeability structure of fluvial-deltaic sandstones, Ferron Sandstone, east-central Utah, in Flores, R.M., ed., Mesozoic of the Western Interior: Society for Sedimentary Geology (SEPM) Guidebook, p. 5–21. Heffern, E.L., and Coates, D.A., 2004, Geologic history of natural coal-bed fires, Powder River basin, USA, in Stracher, G.B., ed., Coal fires burning around the world: a global catastrophe: International Journal of Coal Geology, v. 59, no. 1–2, p. 25–47. Hintze, L.F., 1988, Geologic History of Utah: Provo, Utah, Brigham Young University, Geologic Studies Special Publication 7, 202 p. Johnson, T.M., Liu, F., and Newfarmer, R.S., 1997, Clear water, blue skies: China’s environment in the new century: Washington, D.C., World Bank, 114 p. Lupton, C.T., 1916, Geology and coal resources of Castle Valley in Carbon, Emery, and Sevier counties, Utah: U.S. Geological Survey Bulletin 628, 88 p. Montgomery, S.L., Tabet, D.E., and Barker, C.E., 2001, Upper Cretaceous Ferron sandstone: Major coalbed methane play in central Utah: American Association of Petroleum Geologists Bulletin, v. 85, no. 2, p. 199–219. Montgomery, S.L., Tabet, D.E., and Barker, C.E., 2004, Coalbed gas in the Ferron Sandstone Member of the Mancos Shale: A major Upper Cretaceous play in Central Utah, in Chidsey, T.C., Jr., Adams, R.D., and Morris T.H., eds., American Association of Petroleum Geologists Studies in Geology 50: American Association of Petroleum Geologists, p. 501–528. National Mining Association, 2005a, U.S. coal production—2003: Washington, D.C., National Mining Association, http://www.nma.org/pdf/c_production_2003.pdf (accessed 26 Aug. 2005). National Mining Association, 2005b, Facts about coal (coal use by state): Washington, D.C., National Mining Association, http://www.nma.org/ statistics/pub_fast_facts.asp (accessed 26 Aug. 2005). Nichols, M.W., Jr., 2001, Report of investigation—underground coal mine explosions, July 31–August 1, 2000, Willow Creek Mine, MSHA Id. No. 42–

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Figure 8. (Stop 6). (A) D.E. Tabet standing next to Ferron No. 7 sandstone collapsed over the actively burning J coal bed. (B) G.B. Stracher using a Pasco thermocouple probe to measure the temperature (106 °C) of smoldering coal in the J bed. (C) G.B. Stracher using a Drager hand pump and tube apparatus to extract coal-fire gas from a borehole into an underground coal mine tunnel at the South Cañon Number 1 Coal Mine fire, Colorado (Stracher et al., 2004). Color changes in reagents in CO2 and CO Drager tubes inserted into the pump permit in situ gas analyses in ppm concentrations. (D) G.B. Stracher using a stainless steel gas canister and extraction line to extract coal-fire gas from the borehole in (C) for chromatographic analysis. The same apparatus in (B), (C), and (D) will be used to demonstrate the measurement and collection procedures to field trip participants at Stop 6.

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02113, Plateau Mining Corporation, Helper, Carbon County, Utah, U.S.: Mine Safety and Health Administration, http://www.msha.gov/minefire/ wcreek2000/willowcreekfinal/wcfinalcvr.htm (accessed 26 Aug. 2005). Office of Surface Mining, 1997, Utah 99 percent of 20-year production from underground mines: Washington, D.C., Office of Surface Mining, http: //www.osmre.gov/pdf/utah.pdf (accessed 26 Aug. 2005). Office of Surface Mining, 2004, Utah awarded $1.7 million to regulate coal mine reclamation: Washington, D.C., Office of Surface Mining, http: //www.osmre.gov/news/062904.htm (accessed 26 Aug. 2005). Powell, A.K., 1994, The United Mine Workers of America, in Powell, A.K., ed., Utah History Encyclopedia: Salt Lake City, University of Utah Press, p. 575–576, http://www.media.utah.edu/UHE/u/ UNITEDMINEWORKERS.html (accessed 25 Feb. 2005). Quick, J.C., Tabet, D.E., Hucka, B.P., and Wakefield, S.I., 2004, The available coal resource for eight 7.5-minute quadrangles in the southern Emery coalfield, Emery and Sevier Counties, Utah: Salt Lake City, Utah Geological Survey, Special Study 112, 37 p. Roosevelt, F.D., 1938, The public papers and addresses of Franklin D. Roosevelt, with a special introduction and explanatory notes by President Roosevelt, v. 4, The court disapproves, 1935: New York, Random House, 686 p., http://www.fdrlibrary.marist.edu/odnlra.html (accessed 26 Aug. 2005). Ryer, T.A., 1981, Deltaic coals of Ferron Sandstone Member of Mancos Shale—predictive model for Cretaceous coal-bearing strata of Western Interior: American Association of Petroleum Geologists Bulletin, v. 65, p. 2323–2340. Ryer, T.A., 1991, Stratigraphy, facies, and depositional history of the Ferron Sandstone near Emery, Utah, in Chidsey, T.C., Jr., ed., Geology of eastcentral Utah: Utah Geological Association Publication 19, p. 45–54. Schwans, P., and Campion, K.M., 1997, Sequence architecture and stacking patterns in the Cretaceous foreland basin, Utah: tectonism versus eustasy, in Link, P.K., and Kowallis, B.J., eds., Mesozoic to Recent Geology of Utah: Provo, Utah, Brigham Young University Geology Studies, v. 42, part 2, p. 105–125. Stracher, G.B., 2002, Coal fires: A burning global recipe for catastrophe: Geotimes, v. 47, no. 10, p. 36–37, 66.

Stracher, G.B., 2003, Coal mine fire—gas and condensation products: Collection techniques for laboratory analysis: Energeia, Center for Applied Energy Research, University of Kentucky, v. 14, no. 5, p. 2 and 4–5. Stracher, G.B., and Taylor, T.P., 2004, Coal fires burning out of control around the world: Thermodynamic recipe for environmental catastrophe, in Stracher, G.B., ed., Coal fires burning around the world: A global catastrophe: International Journal of Coal Geology, v. 59, no. 1–2, p. 7–17. Stracher, G.B., Renner, S., Colaizzi, G., and Taylor, T.P., 2004, The South Cañon Number 1 Coal Mine fire: Glenwood Springs, Colorado, in Nelson, E.P., and Erslev, E.A., eds., Field trips in the southern Rocky Mountains, USA: Geological Society of America Field Guide 5, p. 143–150. Tabet, D.E., Hucka, B.P., and Sommer, S.N., 1995, Maps of total Ferron coal, depth to the top, and vitrinite reflectance for the Ferron Sandstone Member of the Mancos Shale, central Utah: Salt Lake City, Utah Geological Survey Open File Report 329, 3 plates, scale 1:250,000. Tabet, D.E., Quick, J.C., Hucka, B.P., and Hanson, J.A., 1998, Available coal resources for the Northern Wasatch Plateau coalfield, Carbon and Emery Counties, in Stringfellow, J., ed., Survey Notes: Salt Lake City, Utah Geological Survey, v. 31, no. 1, p. 1–2. Taniguchi, N.D., 1990, Old king coal: A long, colorful story, in Murphy, M.B., ed., Beehive History, v. 16: Salt Lake City, Utah State Historical Society, p. 14–17, http://historytogo.utah.gov/coal.html (accessed 26 Aug. 2005). Thatcher, L., 1994, Utah State symbols, in Powell, A.K., ed., Utah History Encyclopedia: Salt Lake City, University of Utah Press, p. 601–604. Tripp, C.N., 1989, A hydrocarbon exploration model for the Cretaceous Ferron Sandstone Member of the Mancos Shale and the Dakota Group in the Wasatch Plateau and Castle Valley of east-central Utah, with emphasis on post-1980 subsurface data: Salt Lake City, Utah Geological Survey Open-File Report 160, 81 p. U.S. Department of Energy, 2004, US coal reserves by state and type-2003 (billion short tons): Washington, D.C., U.S. Department of Energy, National Mining Association, http://www.nma.org/pdf/c_reserves.pdf (accessed 26 Aug. 2005). World Resources Institute, 1999, 1998–1999 World Resources: A guide to the global environment, environmental change and human health: New York, Oxford University Press, 369 p.

Printed in the USA

Geological Society of America Field Guide 6 2005

Latest Pleistocene–early Holocene human occupation and paleoenvironmental change in the Bonneville Basin, Utah–Nevada David Rhode Desert Research Institute, 2215 Raggio Parkway, Reno, Nevada 89512, USA Ted Goebel Department of Anthropology, University of Nevada, Reno, Nevada 89557, USA Kelly E. Graf Department of Anthropology, University of Nevada, Reno, Nevada 89557, USA Bryan S. Hockett Bureau of Land Management Elko Field Office, 3900 Idaho Street, Elko, Nevada 89801, USA Kevin T. Jones Antiquities Section, Utah Division of State History, 300 Rio Grande, Salt Lake City, Utah 84101, USA David B. Madsen Texas Archaeological Research Laboratory, J.J. Pickle Research Campus, The University of Texas at Austin, 10100 Burnet Road, Austin, Texas 78758, USA Charles G. Oviatt Department of Geology, Kansas State University, Manhattan, Kansas 66506, USA Dave N. Schmitt Desert Research Institute, 2215 Raggio Parkway, Reno, Nevada 89512, USA

ABSTRACT On this field trip, you will visit two important archaeological cave sites that provide the most compelling evidence for latest Pleistocene and earliest Holocene human occupation in the Bonneville Basin. Danger Cave, located near Wendover, Utah/Nevada, is famed for its deeply stratified archaeological deposits dating as old as 10,300 radiocarbon yr B.P., when the remnant of Lake Bonneville stood at the Gilbert shoreline. Bonneville Estates Rockshelter, located south of Danger Cave at the Lake Bonneville highstand shoreline, also contains well-preserved stratified deposits, including artifacts and cultural features dated to at least 11,000 radiocarbon yr B.P., making it one of the oldest known archaeological occupations in the Great Basin. We describe results of our recent research at these sites and show the stratigraphic evidence for these earliest human occupations. We also review recent work at the Old River Bed Delta, on Dugway Rhode, D., Goebel, T., Graf, K.E., Hockett, B.S., Jones, K.T., Madsen, D.B., Oviatt, C.G., and Schmitt, D.N., 2005, Latest Pleistocene–early Holocene human occupation and paleoenvironmental change in the Bonneville Basin, Utah–Nevada, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 211–230, doi: 10.1130/2005.fld006(10). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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D. Rhode et al. Proving Ground, that has documented hundreds of Paleoarchaic occupation sites dating 11,000–8500 radiocarbon yr B.P. Together these localities give us an unparalleled picture of human occupation during the first few thousand years of known human occupation in the region, during a time of dramatic environmental change. Packrat middens, pollen sampling localities, and geomorphic features that illustrate the history of Pleistocene Lake Bonneville and the environmental history of the western Bonneville Basin will also be observed on this trip. Keywords: archaeology, Late Pleistocene, early Holocene, Bonneville Basin, Bonneville Estates Rockshelter, Danger Cave, Old River Bed, Paleoindian.

INTRODUCTION Recent investigations into latest Pleistocene–early Holocene human occupation of the Bonneville basin, northwest Utah and adjacent Nevada, shed considerable new light on the nature of earliest human adaptations in the context of dramatic environmental changes in the region. In this field guide, we present results of recent archaeological work at the Old River Bed paleodelta, Danger Cave, and Bonneville Estates Rockshelter (Fig. 1). This guide departs from the traditional “road-log” style because the stops described herein are either inaccessible (for government security reasons) or require escorted permission to visit. All historic properties on federal lands, including Bonneville Estates Rockshelter and those in the Old River Bed paleodelta, are protected under the National Historic Preservation Act and the Archaeological Resources Protection Act. Danger Cave is secured under lock and key as a State Park. General directions to the locations of these stops are provided below, but access to them must be arranged in advance with the appropriate land manager. PALEOENVIRONMENTAL CONTEXT

Figure 1. Extent of Bonneville Basin, northwest Utah and adjacent Nevada and Idaho, showing locations of field trip stops and other localities mentioned in the text. BER—Bonneville Estates Rockshelter; BL—Blue Lake; DC—Danger Cave; HC—Homestead Cave; LG—Lake Gunnison; ORB—Old River Bed delta sites; ORBT—Old River Bed Threshold; PSG—Public Shooting Ground; RRP—Red Rock Pass; SLC—Salt Lake City (City Creek locality).

People first occupied and settled into the Bonneville basin during a time of great transition in the character and operation of geomorphic and hydrologic systems, in the content and distribution of vegetation associations, and in the composition and abundance of faunas. This section briefly introduces the environmental context for initial human occupation from ca. 15,000–7500 radiocarbon yr B.P. About 15,000 radiocarbon yr B.P., when Pleistocene Lake Bonneville reached its highstand, it covered 51,700 km2 of northwestern Utah (Fig. 1) to a maximum depth of ~372 m. At this level (1552 m, adjusted for isostatic rebound; Fig. 2), the lake overflowed into the Snake River drainage through a natural alluvial divide at Zenda, Idaho. This huge, cold lake was supported by very low postglacial temperatures relative to today, combined with moderately enhanced precipitation resulting from a southward shift of the mean jet stream (Thompson et al., 1993). Evidence from packrat middens (Rhode and Madsen, 1995; Rhode, 2000a; Thompson, 1990) indicates that, when Lake Bonneville filled much of the lowlands, the mountains west of the Bonneville basin were covered with a subalpine parkland

Latest Pleistocene–early Holocene human occupation and paleoenvironmental change dominated by sagebrush, with scattered stands of spruce, smaller amounts of limber pine, and shrubs such as currant and snowberry as common associates. The presence of mesophilic shrubs and grasses, along with the apparent dominance of spruce over limber pine, suggests the period was cold and relatively moist. Fossil remains of musk ox, mountain sheep, mammoth, horse, camel, bison, mastodon, short-faced and black bear, ground sloth, peccary, and other large beasts document the presence of a diverse megafaunal community. Around 14,500 radiocarbon yr B.P., the unconsolidated alluvial dam at Zenda collapsed catastrophically (Fig. 2), producing a massive flood of 4750 km3 of water into the Snake River drainage, with a calculated peak discharge of ~106 m3/s, roughly equivalent to the mean discharge of all the world’s rivers combined (Jarrett and Malde, 1987; O’Connor, 1993). Within a year, the spill reached resistant bedrock at Red Rock Pass, Idaho, and the lake stabilized at the Provo shoreline, ~1444 m, where it remained for at least several centuries. Lake Bonneville began to recede from the Provo level sometime after 14,000 radiocarbon yr B.P. This regressive phase lasted for the next few thousand years with several fluctuations, but its tempo is the subject of much current debate and ongoing research (Fig. 2). Oviatt (1997; Oviatt et al., 1992) proposed that the lake’s decline began ca. 14,000 radiocarbon yr B.P. (see also Sack, 1999), gradually at first and then more rapidly after ca. 12,500 radiocarbon yr B.P., reaching the level of modern Great Salt Lake by 12,000 radiocarbon yr B.P. More recently, Oviatt et al. (2005) suggest that the decline from the Provo shoreline began later, ca. 13,000 radiocarbon yr B.P., declining gradually until ~12,000 radiocarbon yr B.P., after which it dropped more rapidly to the level of Great Salt Lake at ca. 11,200 radiocarbon yr B.P. An even later age of regression is proposed by Godsey et al. (2005), who examined a large suite of dates and geomorphic profiles from the Provo shoreline complex. They concluded that the lake fluctuated significantly from 14,000–12,500 radiocarbon yr B.P. but that it existed at the Provo shoreline as late as 12,000 radiocarbon yr B.P., after which it declined precipitously to low levels by ca. 11,500 radiocarbon yr B.P. The reconstruction posited by Godsey and colleagues may conflict with other evidence such as the date of initial overflow of Lake Gunnison into the Great Salt Lake Desert (Fig. 2). The different scenarios have implications for the relationship of Lake Bonneville’s decline to global climatic forcing at the end of the Pleistocene, as well as more local considerations such as the antiquity of delta and fluvial channel deposits and the development of wetlands in places such as the Old River Bed (see Stop 1). Fish remains from Homestead Cave, in the Lakeside Range (Fig. 1), provide constraints on the timing of Lake Bonneville’s demise (Broughton, 2000; Broughton et al., 2000). Abundant remains of eleven species of cold- and freshwater adapted fish, including bull and cutthroat trout, whitefish, Bonneville cisco, and sculpin, were found in deposits dating to ca. 11,300–10,400 radiocarbon yr B.P. These remains signal catastrophic die-offs of the native coldwater fishes as the lake receded and became

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warmer and/or more saline (Broughton et al., 2000). The Homestead record suggests the lake was sufficiently cold and fresh to support these fishes prior to ca. 11,300 radiocarbon yr B.P., but by then had declined. Strontium-isotope ratios of fish bones from Homestead Cave and other sites support the idea of a high freshwater lake prior to ~12,000 radiocarbon yr B.P., but a much lower lake by ca. 11,200 radiocarbon yr B.P. (Broughton et al., 2000; Quade, 2000). As the lake declined, the mesophilic parkland that characterized the vegetation in the western Bonneville basin was replaced by a limber pine–sagebrush mosaic, at least at lower elevations above the Bonneville shoreline (Rhode, 2000a; Rhode and Madsen, 1995). This transition at ca. 13,200 radiocarbon yr B.P. suggests a significant drying trend within a still-cool temperature regime. Limber pine became the widespread and dominant tree species, while spruce and mesophilic shrubs largely dropped from the midden record. Common juniper and snowberry were frequent associates. The abundance of limber pine at lower elevations implies the climate was substantially cooler than today (6–7 °C during the growing season), with precipitation slightly greater than modern levels (Rhode and Madsen, 1995). This limber pine–sagebrush association persisted until at least 11,800 radiocarbon yr B.P. in the lower elevations around the Bonneville shoreline. The decline of Lake Bonneville to levels at least as low as 1280 m by 11,200 radiocarbon yr B.P. signals the end of the Bonneville cycle and the beginning of the Great Salt Lake cycle (Fig. 2; Oviatt et al., 1992, 2005). It correlates with a brief period that Haynes (1991) termed the “Clovis drought.” This sharp drying episode might have lasted only a few centuries, but it may have been responsible for significant vegetation changes and the extirpation of Pleistocene megafauna in the Bonneville basin (Madsen, 2000, p. 171), as well as the depletion of the Bonneville fish fauna. The limber pine–sagebrush mosaic widespread before 11,800 radiocarbon yr B.P. was replaced by a shrubland dominated by sagebrush and shadscale, lacking conifers, that was in place by at least 11,000 radiocarbon yr B.P. After that time,

Figure 2. Alternative scenarios for timing of regression of Lake Bonneville from highstand. A—Oviatt et al., 1992, Oviatt, 1997. B—Oviatt et al., 2003, 2005. C—Godsey et al., 2005. GSL—Great Salt Lake.

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limber pine retreated to mid-montane settings and may have persisted in small locally protected pockets at lower elevations. The decline of the lake allowed marshes to develop on the old lakebed in various parts of the Basin (Oviatt et al., 2003). Marshes and wetlands are thought to have developed along the channels of the Old River Bed delta (Stop 1), at Public Shooting Grounds (Oviatt et al., 2005), and at Fish Springs, where peats as old as 11,400 radiocarbon yr B.P. have been found (Godsey et al., 2005). These wetlands may have served as magnets for occupation by some of the first human inhabitants of the region. While the timing of the regressive phase is still uncertain, a subsequent transgressive phase is somewhat better understood (Oviatt et al., 2005). This transgression resulted in the formation of the Gilbert shoreline (Eardley et al., 1957; Currey, 1982), with an elevation ranging between 1293 and 1311 m. The variation reflects isostatic adjustments to some degree (Currey, 1980), but may be the result of other factors such as wave energy, sediment supply, slope morphology, etc. (cf. Oviatt et al., 2005). Currey (1990) and Benson et al. (1992) thought that the transgression to the Gilbert shoreline began by ca. 12,000 radiocarbon yr B.P., progressing through four successively higher transgressive stages to reach the uppermost Gilbert shoreline ca. 10,300 radiocarbon yr B.P. That timing seems to conflict with the more recent estimates of the Bonneville regression, however. Work by Oviatt et al. (2005) at the Public Shooting Grounds, northeast Great Salt Lake (Fig. 1), indicates the Gilbert shoreline deposits there date between 10,500–10,000 radiocarbon yr B.P. (12,900– 11,200 cal B.P.). This transgressive phase roughly coincides with the Younger Dryas cold event. Oviatt et al. (2005) caution that the two events, the Younger Dryas and the lake rise to the Gilbert shoreline, presently appear to be only partially coincident in time; a causal connection needs stronger confirmation. Broughton (2000) suggests that stocks of the Bonneville fish fauna, adapted to cold fresh water, may have rebounded as the lake rose to the Gilbert Shoreline, but by 10,400 radiocarbon yr B.P. most of those fish stocks were apparently decimated as well (except perhaps the more salinity- and warmth-tolerant Utah chub, which can survive in springs and marshes). Sometime around 10,000 radiocarbon yr B.P., the Gilbert transgression waned and the lake returned to low levels once again. Extensive wetlands expanded in several parts of the Bonneville basin lowlands: for example, along the Old River Bed, where wetlands had covered the lowlands since before the Gilbert transgression (Oviatt et al., 2003); at the Public Shooting Grounds (Oviatt et al., 2005); along the margins of the Silver Island Range, near Danger Cave where peats located at an elevation slightly lower than the Gilbert Shoreline date to 9450 ± 150 radiocarbon yr B.P.; and at Blue Lake near Bonneville Estates Rockshelter (Fig. 1), where basal peat deposits date to 9590 ± 50 radiocarbon yr B.P. (Beta-197282). Elsewhere, an exposure of the City Creek fan east of Great Salt Lake (Fig. 1) revealed a series of at least 12 streamside, gallery-forest deposits interbedded with alluvial sand and gravel. Wood from two of these forest-floor mats produced radiocarbon dates of 9670 ± 80 radiocarbon yr

B.P. (Beta-145038) and 9360 ± 60 radiocarbon yr B.P. (Beta142291), respectively. These ages indicate that gallery forests grew here at an altitude only slightly higher than the long-term modern average altitude of Great Salt Lake (GSL) (1280 m), at a time when isostatic rebound was still under way. The Holocene Great Salt Lake probably oscillated significantly in elevation and surface area. Between 10,000 and 9000 radiocarbon yr B.P., the geographic center of the lake was located west of its current center, and its surface area was much larger than at present (Bills et al., 2002). A presumed early Holocene shoreline at ca. 1290 m (Murchison, 1989) may date ca. 9700 radiocarbon yr B.P. (Murchison and Mulvey, 2000), but this lake rise and its age are both uncertain. Early Holocene vegetation in lowlands of the western Bonneville basin was dominated by xerophilic shrubs such as sagebrush, shadscale, greasewood, and horsebrush, but also included hackberry in rocky settings. Hackberry prefers more summer precipitation than is typically available today, so summers were likely to have been somewhat moister than now (Rhode, 2000b). Upland woodlands were dominated by Rocky Mountain juniper. These shrublands and woodlands probably reflect a somewhat cooler and more mesic environment than exists today, with greater sagebrush abundance than now (Rhode, 2000b). Well-dated faunal sequences from Homestead Cave and Camelsback Cave (near the Old River Bed delta; Fig. 1) indicate more mesic and cooler temperatures as well (Grayson, 1998; Madsen, 2000; Schmitt et al., 2002). Under these slightly more mesic conditions, a larger early Holocene Great Salt Lake was probably supported, even though it might not have greatly exceeded the elevation of the lake today. This is because the basin configuration differed from today. The Eardley threshold north of the Lakeside Mountains had not yet formed as a result of isostatic rebound, and the Great Salt Lake occupied a much larger area in the western Great Salt Lake Desert. The existence of this large but shallow lake would have been possible only under conditions of less evapotranspiration than today, hence slightly cooler or moister conditions as reflected in the vegetation, fauna, and well-watered marshlands. By ca. 8500 radiocarbon yr B.P., increased aridification had resulted in a more open shrubland increasingly dominated by shadscale and other more drought-tolerant plants (Bright, 1966; Beiswenger, 1991; Thompson, 1992). The extraordinarily rich faunal record from Homestead Cave clearly demonstrates how increasing aridity and vegetation changes during the early Holocene dramatically affected the relative abundance of a wide range of small mammals, including cottontails, pygmy rabbits, kangaroo rats, voles, pocket and harvest mice, pocket gophers, woodrats, and marmots (Grayson, 1998, 2000; Madsen, 2000). By 8000 radiocarbon yr B.P., lowlands in the Bonneville basin had reached a character much more like that of today (Grayson, 2000; Rhode, 2000b; Schmitt et al., 2002). Warmth-tolerant conifers such as singleleaf piñon and Utah juniper migrated into the region and established montane woodlands by ca. 6500 radiocarbon yr B.P., and the transition to a modern regional ecosystem was essentially complete.

Latest Pleistocene–early Holocene human occupation and paleoenvironmental change The first people to inhabit the Bonneville basin saw the region when it was still very different from the modern. Present evidence demonstrates that humans began to occupy the region by ca. 11,000 radiocarbon yr B.P. and possibly somewhat before. The record of these initial occupations, and how people subsequently adjusted to latest Pleistocene and early Holocene environmental shifts, is best illustrated by the three archaeological sites or site complexes discussed or visited on this field trip: the Old River Bed sites, Danger Cave, and Bonneville Estates Rockshelter. STOP 1. OLD RIVER BED PALEODELTA OPEN SITES The Old River Bed paleodelta and wetlands are located on the southeast side of the Great Salt Lake Desert (Fig. 1), south of Interstate 80 (I-80) between mile posts 10 and 40, and north of the Simpson Springs–Callao road. The area is generally inaccessible as it is almost wholly contained within lands under the control of the U.S. Air Force Utah Test and Training Range and the U.S. Army Dugway Proving Ground, which lies at the north end of the Old River Bed. We will not visit this stop on the field trip, but we will observe the area from a distance and discuss the sites and setting while we are at Stop 3, Bonneville Estates Rockshelter. For those who wish to visit the Old River Bed and the extreme southern portion of the paleodelta on public lands, take the Tooele exit on I-80 and travel southbound on State Highway 36 for 42 miles (68 km) to Vernon, then turn west on the Simpson Springs–Callao Road (the old Pony Express Route) and drive ~35 mi to the Old River Bed. The Old River Bed is an abandoned river valley eroded into deposits of Lake Bonneville in western Utah (Fig. 1). During the regressive phase of Lake Bonneville, a shallow lake in the Sevier basin overflowed to the north. The river created by this overflow eroded a meandering, narrow valley in the fine-grained lake sediments on the basin floor (Oviatt et al., 2003). Sometime after ca. 9000 radiocarbon yr B.P., water ceased to flow in the Old River Bed and environmental conditions along the channel began to approach those found at present. During the roughly 3000 yr of its existence, however, the water in the river fed a large marsh–wetland system at the Old River Bed delta and supported a riverine environment along its length. This 3000 yr interval corresponds almost exactly to the earliest Paleoarchaic phase of human occupation in the Bonneville basin (Beck and Jones, 1997). Foragers have been drawn to these rich marsh–wetland ecosystems throughout the human history of the Great Basin (Madsen, 2002), but sites of the Paleoarchaic period are particularly associated with Great Basin wetlands. Most major wetlands in the Great Basin lie at the end of major river systems, such as the Humboldt, Bear, and Carson Rivers, and have been in existence for at least the period of human occupation in the region. As a result of both continuous use of these marshes by foragers and erosional-depositional cycles associated with Holocene climatic changes, intact Paleoarchaic sites are relatively rare. The Old River Bed delta differs

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in that it was forming for only a limited time during the Paleoarchaic period, and, while erosion has taken place since that time, there has been no subsequent lateral migration of streams that would result in the disturbance of early sites. After ca. 8500 radiocarbon yr B.P., when the wetlands dried up, the area became unattractive to hunter-gatherers, so subsequent human disturbance has been minimal. Primary Geomorphic Features of the Old River Bed This section briefly summarizes the geomorphic features of the Old River Bed wetland, described elsewhere in greater detail (Oviatt et al., 2003). The wetlands of the Old River Bed delta once covered ~700 km2 in the Great Salt Lake Desert. An abrupt boundary in the delta separates the present-day groundwater-discharge mudflats from well-drained, fine-grained sheetflow and eolian deposits (Fig. 3). We informally use “gravel channel” and “sand channel” for fluvial landforms and deposits on the mudflats (Oviatt et al., 2003, their Figures 3 and 4). Gravel channels are deposits of coarse sand and gravel that in planview are straight to curved and digitate, and have abrupt bulbous ends. In transverse cross section, gravel channels are topographically inverted, with the crests of the gravel deposits standing one to four meters

Figure 3. Location of Old River Bed delta sites on Dugway Proving Ground, as shown in Figure 1. Note the gravel channels on the mudflats (dark anastomosing features on the mudflats), sand channels (lighter sinuous features around gravel channels), and the abrupt boundary between the mudflats and the well-drained, fine-grained deposits on the flat desert floor. Gravel and sand channel locations that have been surveyed for archaeological sites are outlined in white, and the locations of known Paleoarchaic sites are marked with black dots. Given the site density in the surveyed areas, there may be as many as 500 Paleoarchaic sites in the delta.

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higher than the surrounding mudflats. They can be traced down to an altitude of 1299 m, where they end abruptly. Gravel channel characteristics indicate the Old River Bed river emptied into a shallow Lake Bonneville (Fig. 4). Lake Bonneville had dropped to approximately the level of the Old River Bed threshold (~1390 m) by ca. 12,500 radiocarbon yr B.P. (Oviatt, 1988), and the gravel channels were produced by river discharge from the Sevier basin across the threshold after this time (Fig. 4). They ceased to form after the lake declined to below 1299 m, probably about the time of the massive fish dieoff around 11,200 radiocarbon yr B.P. (Broughton et al., 2000). Sand channels are found in the same general area of the mudflats as the gravel channels (Fig. 3). They are less topographically inverted and are truncated by the mudflat surface. Where they have been protected, they may stand as much as 1.2 m above the surrounding mudflats, and in other areas they typically stand

~0.5 m above the mudflats. Sand channels are not easy to identify on the ground due to deflation of the mudflats, but from the air the preserved sand channel roots exhibit well-developed meander-scroll patterns. Sediments in sand channels consist of fine to coarse cross-bedded sand, and locally include mud containing organic mats, mollusks and bones of Utah chub, a fish adapted to warm and slightly saline waters. Most sand channels end at an altitude between 1301 and 1303 m. Intermediate channel forms, which are straighter and smaller in width than sand channels and locally contain some gravel, can be traced to altitudes as low as 1285 m in the west-central Great Salt Lake Desert. Sand channels are younger than gravel channels. The scoured bottoms of sand channels are topographically lower than the bottoms of gravel channels, and sand-channel deposits are inset into gravel-channel deposits. Sand channels were produced by perennial rivers that were active during the period from at least 11,000–8800 radiocarbon yr B.P. (see Oviatt et al., 2003, their Table 1). Sometime prior to 11,000 radiocarbon yr B.P. the Sevier basin stopped overflowing, and stream flow in the river was reduced, though still substantial enough to carry coarse sands in channels. Water in these sand channels fed a large wetland-marsh ecosystem over much of what is now the Great Salt Lake Desert. During this episode, exposed underflow fan and lacustrine deposits began to deflate in what are now mudflats, partially exhuming the gravel channels. Human Occupation of the Old River Bed Wetlands

Figure 4. Schematic diagrams showing changing lake level during the regression of Lake Bonneville and its effect on river flow in the Old River Bed valley and on the deposition of gravel channels at Dugway Proving Ground. (A) Early regressive phase of Lake Bonneville. Lake surface is above the Old River Bed threshold, and endogenic calcium carbonate (marl) is being deposited. (B) Lake has dropped below threshold and river flow has begun in the Old River Bed valley; suspended-load sediments are spread to the north by underflow currents in the regressing lake, and silt and clay are deposited over the marl. (C) Lake level continues to drop, and bedload sediments are deposited over the underflow muds; gravel channels prograde into the shallow lake.

Archaeological sites in the Old River Bed delta are associated with the exposed sand and gravel channels (Fig. 3). The sand channel streams meandered extensively through the delta, creating a vast wetland system with relatively few areas suitable for habitation or for activities not directly associated with foraging in the marsh itself. The only dry areas were probably the partially exposed gravel channels and natural levees along the sand channel margins, and it is these areas where sites appear to be concentrated. We have conducted archaeological inventories of 52 km of the more than 200 linear km of exposed channels within Dugway Proving Ground and located 51 Paleoarchaic archaeological sites directly associated with channels or immediately adjacent wetlands (Fig. 3). An additional five sites not directly associated with channel features have also been identified. This density, together with archaeological investigations at the extreme northwestern toe of the delta (Arkush and Pitblado, 2000; Carter and Young, 2001), suggests as many as 500 or more Paleoarchaic sites may be present within the delta area. These sites consist of surface arrays of a few dozen to hundreds of basalt and obsidian flakes and tools. None of these sites has yet been excavated, and they are not directly dated. But their ages can be estimated by their relationship to channel features and by typological dating of associated diagnostic artifacts. All the sites postdate formation of the gravel channels, and, hence, date to the period between 11,000 and 9000 radiocarbon yr B.P. (Oviatt et al., 2003).

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Site Types and Tool Forms

Paleoarchaic Mobility

The Paleoarchaic sites in the Old River Bed delta area take one of two principal forms. The majority are linear features along the margins of topographically inverted gravel channels or adjacent to sand channels. These linear sites have a high ratio of finished tools to flaking debris and most of the tools consist of stemmed, hafted bifaces. Debitage and some tools feather out onto adjacent mudflats that were probably wetlands when the sites were occupied. The other, less common, site type consists of diffuse scatters of lithic debitage and stone tools on mudflat surfaces. These are often hundreds of meters in diameter, and in places merge into a background scatter of cultural materials covering much of the mudflats within the delta area. Tool to debitage ratios in these sites are reduced and they appear to represent resource procurement activities within the wetlands themselves. The sites are characterized by a variety of Great Basin Stemmed bifaces and crescents dating to 11,500–8500 radiocarbon yr B.P. (Fig. 5) (Beck and Jones, 1997). Many tools appear to have been scavenged and reworked from earlier deposits and were repeatedly resharpened to such an extent that the amount of cutting edge above the hafting element is minimal. This is particularly true for a unique form of Great Basin Stemmed biface we term Dugway Stubbies. Large basalt flakes and cores, which appear to have been used as tools, are also common. These cores are usually in the form of large thin bifaces (Fig. 6), but include platform cores used to produce short blades. Although we have no direct dates for any of these tool forms, the sample size is sufficiently large at 23 sites to determine a relative chronology based on seriation of stemmed bifaces and crescents (Fig. 7). Stone used for making these tools consists primarily of local basalts and Topaz obsidian from sources less than 50 km from the Old River Bed delta. Minor amounts of Browns Bench obsidian, from a source near the junction of the Utah-Nevada-Idaho borders, are also present.

Paleoarchaic sites in the Old River Bed delta fit easily within the Western Stemmed Tradition characterized by an adaptation to wetland ecosystems around the many shallow lakes on valley floors in the Great Basin during the PleistoceneHolocene transition (Willig et al., 1988). The kind of mobility pattern characteristic of foragers on the Old River Bed delta, while generally similar to that found among Paleoarchaic foragers elsewhere in the Great Basin, appears to have differed in the Old River Bed delta due to the size of the marsh ecosystem. In most of the small, isolated Great Basin valleys, wetland ecosystems were small relative to the Old River Bed delta, and both theoretical models and limited empirical data suggest foragers employed a high mobility pattern characterized by the use of a variety of widely-spaced toolstone sources, large flake and biface tools, high percentages of scrapers, and a diet narrowly focused on wetland resources (Graf, 2001; Elston and Zeanah, 2002; Huckleberry et al., 2001; Beck et al., 2002; Jones et al., 2003). In the Old River Bed delta area, on the other hand, biface tools are extremely small and often extensively reworked, toolstone sources are local and limited in number (much like later Archaic and Fremont toolstone use in the region) (Schmitt and Madsen, 2005), and scrapers are relatively uncommon (cf. Arkush and Pitblado, 2000), suggesting a more restricted pattern of movement between sites. Much of this difference may be due to resource patch size. Elsewhere in the Great Basin, wetland ecosystem patches are small and resources were quickly depleted, necessitating frequent moves of relatively large distances between foraging patches (Madsen, 2002; Elston and Zeanah, 2002; Jones et al., 2003). The wetland ecosystem on the Old River Bed delta was enormous relative to these other small Paleoarchaic wetland foraging localities, and, while movement within the wetland may have been almost as frequent, the distances involved were much shorter. As a result, movement outside the Old River Bed delta area to other foraging locations only occurred after extended periods of stay within the marsh. In turn, the ability to refurbish tool kits frequently at a variety of widely separated toolstone sources was limited and is reflected in the extensive reworking of tools and the few toolstone types found at Old River Bed delta sites.

Foraging Adaptations We have little direct evidence of the foraging activities of the Paleoarchaic people who occupied the sand channel sites, other than that they likely focused on the marsh-wetland resources that dominated the Old River Bed delta landscape at the time. What is most remarkable about the artifact complex we have identified is that it lacks groundstone, suggesting that seed collecting and processing was not part of the subsistence focus. Whether foraging was limited to large and small game animals or included other plant resources such as marsh rhizomes is presently unknown. In the western Great Basin, Paleoarchaic foragers around the Stillwater marsh area were eating small fish at the time the Old River Bed sand channel sites were occupied (Napton, 1997), and it is possible that the fish in the Old River Bed streams were also being exploited. It now appears that seed grinding was not a significant part of foraging strategies in the Bonneville basin until after ca. 8600 radiocarbon yr B.P., about when the Old River Bed wetlands were finally eliminated (Rhode et al., 2006).

STOP 2. DANGER CAVE Danger Cave is located in the Silver Island Range on the edge of the Great Salt Lake Desert, just northeast of the town of Wendover, Utah, where I-80 crosses the Nevada-Utah line (Fig. 1). Now the centerpiece of a small State Park, it is accessible via dirt road from the Bonneville Speedway exit off I-80. Take Exit 4 northward past the truck stop, then turn left onto a westbound paved road and proceed 2.4 km to a dirt road angling northwest. Proceed on this road 1.3 km to a road heading southwest along flank of Silver Island Range. Head southwest along this dirt road 1.0 km (watch for gullies!) until

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Figure 5. View of typical Great Basin Stemmed bifaces found on Old River Bed delta sites. Note the generally irregular forms, extensive reworking, and weathering on all these tool types.

Latest Pleistocene–early Holocene human occupation and paleoenvironmental change you reach a widened parking area, with the fenced-in cave ~50 m away to the west. The cave entrance is fenced to prevent unauthorized visits. To obtain entry into the cave, contact Utah State Parks or the state archaeologist, Antiquities Section, Utah Division of State History. Danger Cave is justly regarded as one of the most important archaeological sites in the Great Basin. The cave attracted human habitation in an otherwise hostile environment throughout the Holocene. A small spring-fed wetland on the playa margin nearby provided water and wetland resources, and the cave itself gave ample shelter from summer heat and winter cold. Excavated by archaeologists several times since the 1940s, its deep multimillennial stacks of well-preserved cultural strata are the fount of some of the most influential concepts in the human prehistory of western North America. Equally amazing, after all the digging by archaeologists (and generations of looters), Danger Cave still has much to offer Great Basin prehistory. We recently assayed its research potential as it relates to early Holocene occupation, and here we describe our results to date. Danger Cave is a large oval chamber ~20 m wide by 40 m long (Fig. 8), formed in Paleozoic limestone by solution weathering in fractures and subsequently enlarged by spalling of the walls and possibly wave action by Lake Bonneville. Situated at an altitude of 1315 m, it is ~311 m below the Bonneville shoreline, 189 m below the Provo, 18 m above the Gilbert shoreline, and ~20 m above the current playa. A date of 13,250 ± 160 radiocarbon yr B.P. was obtained on outermost tufa within the cave, giving a minimum limiting age of its submergence beneath Lake Bonneville’s waters. Lake Bonneville exited Danger Cave sometime prior to ~11,500 radiocarbon yr B.P., as indicated by dates on uncharred wood and sheep dung (Jennings, 1957).

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Figure 6. Large basalt bifaces from the Old River Bed delta. Thinning flakes from these large bifaces appear to have been a principal tool used by Paleoarchaic foragers in the delta. Overshot flakes similar to that seen on this surface are characteristic of Clovis biface forms (M. Collins, 2005, personal commun.).

History of Investigations The site was first archaeologically sampled in the early 1940s by Elmer Smith, who limited his testing to the very front of the cave, beneath the drip line. Its real claim to fame, however, came with the excavation campaigns led by Jesse Jennings

Figure 7. Seriation of Great Basin Stemmed bifaces from 23 Old River Bed delta sites. Biface types are arrayed left to right, sites are arrayed bottom (oldest) to top (youngest) according to the best fit of the “battleship curves” of the seriation. Width of the bars represents proportion of biface types at each site. Pinto points are generally thought to date after ~9000 radiocarbon yr B.P. The age of the other types is unknown, but the time span represented by this seriation is ~11,000–8800 radiocarbon yr B.P.

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Figure 8. Plan map of Danger Cave, showing excavation areas of various investigators discussed in text. Note position of stratigraphic profiles depicted in other figures.

in 1949–1953. Jennings and his University of Utah field crews removed an impressive block in the front half of the cave, exposing an extraordinary stack of well-preserved primarily culturallyderived sediments (Fig. 9). Jennings used the then-new radiocarbon dating method to determine that these deposits span more than 10,000 yr (Jennings, 1957). That experience led Jennings to devise his Desert Culture concept, which proposed that hunter-gatherer lifeways, similar to those of the historically known Shoshonean peoples (Steward, 1938), had stretched back essentially unchanged for over ten millennia. The Desert Culture concept influenced archaeological thought about hunter-gatherer adaptations continentwide, and stimulated much regional research that refined and ultimately refuted some of its conclusions. Although the Great Basin archaeological record now demonstrates significant variability in foraging lifeways through time and across space, the concept still underlies much current regional thought, in the

form of the “Desert Archaic” concept and derivatives such as “Paleoarchaic.” In 1968, a new generation of Jennings’s students, including Gary Fry, David Madsen, and others, returned to Danger Cave to obtain sediments that might allow them to learn more about changing subsistence practices and paleoenvironmental change. They excavated a trench ~8 m farther back in the cave, beyond the back wall of Jennings’s block excavation (Fig. 8), and collected samples of the strata and other materials including human paleofecal material (coprolites). Partial results of these investigations were later published by Harper and Alder (1972) and Fry (1976). In 1986, a block of well-exposed intact deposits was excavated by David Madsen and associates to address new questions about the occupation history of the cave and use of key dietary plants (Madsen and Rhode, 1990; Rhode and Madsen, 1998). These excavations (Fig. 10) provided fine-grained detail about the occupational sequence not afforded by Jennings’s broader treatment. However, the excavation was necessarily more restricted in spatial scope, and the block’s location near the front of the cave (Fig. 8) limited preservation of vegetal remains and other perishable artifacts to the past 8000 yr. Preserved materials dating earlier than that were buried elsewhere in the cave, but these were not obtained in the 1986 excavation. To find earlier well-preserved deposits in other parts of the cave, the Utah Division of State History and the Division of State Parks authorized a reconnaissance of areas previously excavated by Jennings and Fry, allowing the removal of backfill to assess the extent and research potential of remaining intact deposits. The re-exposure of intact deposits serves to enhance the educational value of the site as a historical component of the State Parks system. It is these efforts, conducted since 2001, that we now describe. The 143 Face and Back Trench Jennings and his students terminated their excavations in 1953 at what they called the 143 face; that is, an east-west sidewall running perpendicular to the 143-ft point on the main grid north-south line (Fig. 8). In our investigations, we were able to expose the original 143 face (at least, the lower third of it) and discovered intact deposits undisturbed by decades of looters. These deposits lay protected beneath large rocks and an impenetrable plate of calcium-cemented ash that had formed from water seepage and cementation. These lower strata contain a remarkably well-preserved set of cultural deposits dating from ca. 8000 to over 10,000 radiocarbon yr B.P. Figure 11 shows a profile drawing of the wall as it appears today, together with an inset of the original profile of the 143 face as depicted by Jennings (1957; Figure 54 therein). The back trench, excavated in 1968, exposed layers of pure to nearly pure pickleweed chaff, the byproduct of processing pickleweed for its edible seeds, in a context that was thought to date to ca. 9000–10,000 radiocarbon yr B.P. (Harper and Alder, 1972). To verify the antiquity of pickleweed processing at Danger

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Cave, we re-exposed this trench area, revealing a very well-stratified set of deposits (Fig. 12). The age of the lowest pickleweed processing layer, found at the base of Stratum 04-11 in Figure 12, is 8570 ± 40 radiocarbon yr B.P. (Beta-193123). The two profiles (Figs. 11 and 12) show the lowermost of the main stratigraphic divisions given by Jennings in his original report (1957). Jennings divided the 10,000 yr occupation of Danger Cave into five levels, DI through DV, separated by layers of roof spall that implied long hiatuses of occupation. The upper three (DIII–DV) were composed of thick layers of wellpreserved vegetal debris, dust, ash, and artifacts, dating from ca. 8000 radiocarbon yr B.P. to historic times. They are largely destroyed in the vicinity of the 143 face, but a remnant still exists at the top of the back trench. Of greatest significance here are the lowest two of Jennings’s levels, which he called DI and DII. The DI Occupation: 10,300 Radiocarbon yr B.P. The DI level holds evidence of the earliest human occupation at the site (Fig. 11). Several small firehearths accompanied by a sparse scatter of artifacts and ecofacts lay on a bed of beach sand (“Sand 1”), capped by a variably thick layer of winddeposited sand containing abundant artiodactyl pellets (“Sand 2”). Jennings’s original radiocarbon dating together with recent radiocarbon dates we obtained from ash lenses exposed in the 143 face, confirm that this earliest occupation took place at 10,300 radiocarbon yr B.P. (Fig. 11). This occupation was likely coeval with the existence of a large shallow lake that covered the Great Salt Lake Desert at the level of the Gilbert Shoreline, just below the mouth of the cave. A small but interesting collection of artifacts was obtained from the DI level. These include one lanceolate projectile point of the Agate Basin style, several unifacial scrapers, a few pieces of possible groundstone artifacts (including stones for grinding ochre), numerous chert and obsidian flakes, several knotted pieces of twine of unknown function, modified bone, and assorted fragments of “food bones,” as Jennings called them. Six human coprolites had been found on this level (Fry, 1976), but they date to later times (Rhode et al., 2006). We took several small samples of sediments from the 143 face, and these contained an abundance of chipped stone waste flakes, though no tools, as well as knotted string fragments and small wood whittling curls. Given the small size of our samples, it is likely that the remaining DI occupation deposit still contains an abundance of artifacts.

Figure 9. The back wall (143 face) of excavations in Danger Cave led by Jesse Jennings in 1949–1953. The deposits consist of extensive beds of vegetal material, mostly pickleweed processing residue, some of it burned, resulting in white ash beds. The area recently exposed is to the right of the photographer.

Figure 10. David Madsen (with shovel) and crew member exposing an intact block of sediments in 1986. Over 106 individual strata were mapped in this block (note tags in wall) and were carefully removed in 36 separate excavatable units for subsequent lab analysis.

The DII Occupations: 10,100–7500 Radiocarbon yr B.P. The overlying DII level is a thick deposit of organic debris, cemented ash, rockspall, bat guano, and artifacts that combines three distinct stratigraphic layers (Figs. 11 and 12). The DII level was initially thought to date between ca. 10,000–9000 radiocarbon yr B.P. (Grayson, 1988, 1993; Jennings 1957, 1978), but our excavation and radiocarbon dating of the 143 face and back trench, together with a reanalysis of Jennings’s original

field notes and radiocarbon dates, now show that the three main layers of DII span a longer duration: an upper layer (called F30 in Jennings’s field notes) dating 8200–7500 radiocarbon yr B.P.; a middle layer (F16) estimated to date ca. 8600–8400 radiocarbon yr B.P.; and a lower layer (F31) resting on DI sands dated ca. 10,100–9800 radiocarbon yr B.P.

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Figure 11. Danger Cave 143 face stratigraphic profile as mapped by Jennings in 1953 (A), and as it appears today (B). Radiocarbon dates were obtained from charcoal-ash stains (F111 and F112), charred vegetation from the lowest part of DII (F115), and pickleweed chaff from the uppermost part of DII (F119).

Figure 12. Stratigraphic profile of intact sediments exposed in the Back Trench area seen on field trip. Strata 04-1 through 04-9 correspond to Jennings’s DIII cultural level, strata 04-10 through 04-13 correspond to the DII level, and stratum 04-14 (base of profile) is the DI level. Height of profile at 107.5 N is 115 cm. Radiocarbon dates were obtained from stratum 04-10 and stratum 04-11.

Latest Pleistocene–early Holocene human occupation and paleoenvironmental change The DII layer contained a much richer and more abundant collection of artifacts than the DI occupation. Jennings (1957) reported the following artifact categories: projectile points, bifaces of various stages, end scrapers and flake scrapers, drills and gravers, choppers and scraper planes, shale knives, utilized flakes and blades, abundant millingstones, fragments of cordage, leather, and basketry, and pieces of ochre and mica. Direct evidence of diet included 11 paleofecal specimens (analyzed by Fry, 1976), numerous chewed-up tule wads (“quids”), and hundreds of “food bones.” Although it is difficult to reconstruct artifact counts from available notes, it is clear that most of the artifacts in the DII level came from the two upper layers (F16 and F30) and postdate 8600 radiocarbon yr B.P. Projectile point types commonly found in the DII level generally postdate 8500 radiocarbon yr B.P. (Aikens, 1970; Beck and Jones, 1997). The DII level contained over 160 millingstone and handstone artifacts, almost all coming from layers postdating 8500 radiocarbon yr B.P. (Rhode et al., 2006). Eleven paleofecal samples from DII all contained pickleweed seeds (Fry, 1976), while numerous thin layers of nearly pure pickleweed chaff in DII clearly demonstrate that pickleweed seed winnowing and processing took place in the cave during DII times. Dating of the coprolites and the earliest pickleweed processing layer from the Back Trench shows that pickleweed processing and consumption began ca. 8600 radiocarbon yr B.P. (Rhode et al., 2006). The artifact inventory and character of occupation pre-dating 8600 radiocarbon yr B.P., the lowermost part of DII, is unfortunately poorly known at present, and there may have been a substantial hiatus of occupation between ~10,000 and 8600 yr ago. The advent of pickleweed processing occurred while extensive wetlands like those at the Old River Bed delta were drying up, as the Bonneville basin underwent a period of significant environmental change under increasing Holocene aridification (Madsen et al., 2001). The timing of small-seed use at Danger Cave suggests that people adopted small seeds in their diets in response to broad-scale aridification in the Bonneville basin. In this regard it is of interest to note that other cave sites in the Bonneville basin began to be occupied at about this time or somewhat later, including Hogup Cave (Aikens, 1970) and Camels Back Cave (Schmitt and Madsen, 2005), which provide evidence for a broad-scale adaptation to desert resources.

the high Bonneville shoreline complex of Pleistocene Lake Bonneville, at an elevation of ~1580 m. It is a large, “openmouthed” rockshelter, ~25 m wide and 10 m high at its mouth and as much as 15 m deep, from front to back (Fig. 13). Within the confines of the rockshelter is more than 250 m2 of excavatable surface area (Fig. 14). Although only 30 km apart, the environmental settings of the Bonneville Estates Rockshelter and Danger Cave are significantly different. Given its position on the Lake Bonneville highstand shoreline, the Bonneville Estates Rockshelter would have become open and available for human occupation as much as 3000 yr earlier than Danger Cave. Bonneville Estates Rockshelter is situated 6 km from the nearest source of fresh water (Blue Lake), whereas Danger Cave is only a few hundred meters from a freshwater spring. Today, vegetation in the vicinity of Bonneville Estates Rockshelter is dominated by shadscale, rabbitbrush, and Indian ricegrass, whereas at Danger Cave vegetation is dominated by shadscale, greasewood, pickleweed, and saltbush (Madsen and Rhode, 1990). Vegetation communities at the two sites likely would have been different during the late Pleistocene and early Holocene as well, with limber pine and sagebrush communities persisting longer in the vicinity of the Bonneville Estates Rockshelter (Rhode, 2000a). As a result of these environmental differences, the two sites contain remains of significantly different human activities, and together they have the potential to provide a detailed portrayal of human adaptive change in the western Bonneville basin since initial colonization more than 10,000 radiocarbon yr B.P. Background The Bonneville Estates Rockshelter is one of 13 rockshelters and caves known to exist in the Permian-aged limestones and dolomites of the Lead Mine Hills. Nearly all of these are associated with prominent wave-cut features that correlate with Lake Bonneville’s various shorelines. Bonneville Estates was

STOP 3. BONNEVILLE ESTATES ROCKSHELTER This site is located ~30 km south of Wendover, off Hwy 93. To visit the site, obtain permission from the Elko Field Office, Bureau of Land Management, Elko, Nevada. To access the site, take Nevada State Highway 93 south from Wendover 17.7 mi to the gravel road to Blue Lake, then follow a dirt road from the Blue Lake road southward to the site. Bonneville Estates Rockshelter is located in the Lead Mine Hills of the Goshute Mountains, Elko County, Nevada, ~30 km south of Danger Cave (Fig. 1). The rockshelter is situated along

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Figure 13. View of Bonneville Estates Rockshelter, 2004.

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Figure 14. Map of Bonneville Estates Rockshelter, showing extent of excavations through 2004.

discovered in 1986 by T. Murphy and S. Dondero of the U.S. Department of the Interior Bureau of Land Management (BLM). At the time, the rockshelter was being actively looted. In 1988, P-III Associates, Inc., under the direction of A. Schroedl, conducted test excavations for the BLM in order to discern whether intact cultural deposits still existed in the rockshelter. They found a well-preserved sequence of cultural components that spanned the past 6000 yr of prehistory (Schroedl and Coulam, 1989). At the time, however, they were not able to penetrate deeper into the rockshelter’s deposits due to budgetary constraints. In 2000, our team resumed test excavations in the rockshelter, to learn whether pre-6000 radiocarbon yr B.P. deposits existed, and to determine the rockshelter’s potential for investigating human paleoecology and adaptive change during the late Pleistocene and early Holocene. By the end of 2001, excavations had opened an area of 20 m2, and in two deep tests we exposed cultural components dating to 10,100, 7400, and 7200 radiocarbon yr B.P. (Goebel et al., 2003). In 2003–2005, we further investigated these early components, and so far have opened an area of nearly 60 m2, focusing on two areas referred to as the West and East blocks (Fig. 14). In the West Block, four components spanning from 10,800–9400 radiocarbon yr B.P. have been identified, and below them is even an older stratum with hearth-like features and lithic artifacts that may date to 12,300 radiocarbon yr B.P. In the East Block, two stratigraphically separate components with hearths have been recognized and 14C dated to between 10,600 and 9400 radiocarbon yr B.P. Descriptions of the stratigraphy and cultural remains thus far recovered from the pre-9000 radiocarbon yr B.P. deposits in the two excavations are presented below, to augment our examinations of the exposures on the field trip.

West Block In the western area of Bonneville Estates Rockshelter, we have identified 21 stratigraphic layers in a profile that reaches 280 cm in thickness and spans from ~15,500 radiocarbon yr B.P. to the present (Figs. 15 and 16). The lower 130 cm of this profile are so far culturally sterile. They consist of a thin band of pebblesized gravels (at the base of the profile) thought to represent the 15,500 radiocarbon yr B.P. highstand beach of Lake Bonneville (stratum A21), and an overlying massive silt and rubble deposit (A20) thought to date to between 15,000 and 12,500 radiocarbon yr B.P. (Fig. 15). Among stratum A20 faunal specimens are the only remains of extinct fauna yet found in the rockshelter: a central phalanx of a medium-sized felid (either extinct North American cheetah or cougar) and a large canid patella possibly of dire wolf. The upper 150 cm of the profile consists of a series of 19 discernible strata rich in perishable artifacts and ecofacts as well as hearths, pits, and other cultural features. The early part of this record, spanning from ca. 12,300 to 9400 radiocarbon yr B.P., can be provisionally grouped into three time-stratigraphic zones: (1) late Pleistocene (12,300 radiocarbon yr B.P.), (2) latest Pleistocene (10,800–10,400 and 10,000 radiocarbon yr B.P.), and (3) earliest Holocene (9440 radiocarbon yr B.P.). Late Pleistocene Possibly the oldest cultural remains thus far exposed in Bonneville Estates Rockshelter occur within stratum A19 (Fig. 16). In an area of ~4 m2, we have unearthed an organic-rich layer of silt that contains unequivocal lithic artifacts (20 flakes),

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Figure 15. Stratigraphic profile of the West Block excavation of Bonneville Estates Rockshelter, exposed in 2000–2003.

Figure 16. Stratigraphic profile of the lower strata preserved in the West Block of Bonneville Estates Rockshelter, exposed in 2004. The left side of the profile shown here lies 1 m west of the profile shown in Figure 15.

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floral and faunal remains (including cottontail rabbit, pygmy rabbit, hare, woodrat, pocket gopher, and marmot), and two hearth-like features have been unearthed. One of these (Feature 03.17) has two associated ages of 12,180 ± 60 and 10,640 ± 60 radiocarbon yr B.P., while the other (Feature 04.13b) has four associated ages of 12,270 ± 60, 12,330 ± 40, 12,390 ± 40, and 10,970 ± 60 radiocarbon yr B.P. The 12,400–12,200 14C assays (Fig. 16) were obtained from charred samples having a resinous and mostly non-woody structure, unidentified at the time but now identified as conifer, possibly limber pine. After receiving the four early ages, we dated a sample of unequivocal sagebrush charcoal from each feature; these have yielded the 14C ages of 10,640 and 10,970 radiocarbon yr B.P. The discrepancy between these two sets of ages for stratum A19 is a problem that will be solved with more excavation, detailed macrobotanical analysis, and 14C dating of multiple materials recovered from the features. Latest Pleistocene Two stratigraphically separate cultural components, strata A18b and A18a, have been delineated that date to the very end of the Pleistocene. The lower of the two (A18b) is a 3- to 5-cm-thick organic-rich, ashy stratum. Five hearth features have been excavated from stratum 18b; four of these have produced ages (on charcoal) of 10,405 ± 50, 10,760 ± 70, 10,800 ± 60, 10,690 ± 70, and 10,540 ± 40 radiocarbon yr B.P. (Fig. 16). Faunal remains identified include not only the same leporids and rodents identified in stratum A19, but also numerous specimens of sage grouse, several medium-sized ungulate (including pronghorn) long bone fragments (some with cut marks and others that are burned), and one charred central phalanx of a black bear. Lithic artifacts include 172 flakes and four tools (a finished but unhafted biface [Fig. 17A], two biface fragments, and a retouched flake). Stratum A18a, which lies immediately above A18b, occurs across nearly the entire West Block (Figs. 15 and 16). It is a 5- to 10-cm-thick band of silt with minimal organics that grades from east to west into a richly preserved stratum of organics. Wood charcoal from the single hearth so far excavated in this component yielded three 14C ages averaging 10,090 ± 30 radiocarbon yr B.P. (Goebel et al., 2003). Among faunal remains, sage grouse bones continue to be rather abundant, as are cut and burned ungulate shaft fragments. Additional species include a shaft fragment of pronghorn, an ungulate shaft fragment that is either deer or mountain sheep, and a complete mandible of an ermine, as well as remains of short-eared owl, screech owl, and pintail. Lithic artifacts include 157 flakes and seven tools (three stemmed-point fragments, two side scrapers [Fig. 17E and 17H], one retouched flake, and one possible hammerstone). Perishable materials include six cordage pieces, one small textile fragment, a worked piece of wood, and several knotted feather quills. Earliest Holocene An early Holocene cultural component occurs in Stratum A17b′, a 5–10-cm thick band of organics that is sealed by a massive rock-fall feature (stratum A17b) and a 30-cm thick set

of woodrat midden deposits (PM1 and PM2) (Figs. 15 and 16). So far, we have excavated an area of <6 m2 of this component, exposing one unlined hearth feature, charcoal from which has yielded 14C ages of 9440 ± 50 and 9430 ± 50 radiocarbon yr B.P. (Fig. 16). Associated faunal remains continue to include leporids, rodents, sage grouse, short-eared owl, pronghorn, and cut and burned ungulate long bone fragments. In particular, one complete sage grouse humerus displays numerous stone tool cut marks near its proximal end. Other species identified include bat, horned lizard, and screech owl. Recovered artifacts include a Haskett stemmed-point midsection fragment (Fig. 17G), a Windust stemmed-point basal fragment (Fig. 17C), a retouched flake, 93 waste flakes, and three pieces of cordage. East Block Our study of the eastern area of Bonneville Estates Rockshelter began with the cleaning of a large looters’ pit against the back wall of the rockshelter, and continued with excavation of a 10m2 area adjacent to this pit in order to expose fresh stratigraphic profiles and to investigate early deposits stratigraphically beneath the looters’ fill (Fig. 15). This excavation is still in progress; however, since 2002 we have unearthed two stratigraphically distinct cultural components spanning from the latest Pleistocene through the earliest Holocene (10,500–9400 radiocarbon yr B.P.), as well as nine middle and late Holocene cultural components spanning from 7250 to 80 radiocarbon yr B.P. (Fig. 18). The upper deposits postdating 7250 radiocarbon yr B.P. consist of rich organics and well-preserved features, but the lower deposits are less wellpreserved and occur in loose silt and rubble with little internal structure. In one 1-m2 test pit, we have further excavated to the base of the rockshelter’s Quaternary-aged deposits, unearthing beach gravels lying unconformably upon bedrock at a depth of ~300 cm below the modern surface. Each of the three early cultural components is described briefly below. Latest Pleistocene The basal cultural component in the East Block is stratum 11/12, a deposit of loose and massive silt and rubble (Fig. 18). In an area of ~4 m2, we have unearthed five oval-shaped stains of charcoal and ash that range from 1 to 5 cm in thickness and are underlain by fire-reddened silt. Each of these hearth features has been 14C dated: the first to 10,380 ± 40 radiocarbon yr B.P., the second to 10,030 ± 50 radiocarbon yr B.P., the third to 10,380 ± 55 radiocarbon yr B.P., the fourth to 10,560 ± 50 and 10,050 ± 50 radiocarbon yr B.P., and the fifth to 9990 ± 50 radiocarbon yr B.P. Associated faunal remains include sage grouse and a variety of leporids, as well as one specimen of deer and numerous ungulate long-bone fragments. Lithic artifacts include one midsection fragment of a stemmed point (Fig. 17B) and seven flakes. Earliest Holocene Near the top of stratum 10, ~10–15 cm above the latest Pleistocene component described above, we have exposed two

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hearth features (Fig. 18). One hearth yielded 14C ages of 9520 ± 60 and 9440 ± 75 radiocarbon yr B.P., and the other hearth yielded 14C ages of 9580 ± 40 and 9570 ± 40 radiocarbon yr B.P. Associated faunal remains include sage grouse, deer, at least one cut ungulate long bone fragment, leporids, rodents, and bats. Lithic artifacts include a finished but unhafted stemmed point, an end scraper on a blade (Fig. 17F), and nine flakes. Stratigraphic Trench—First Clovis Find at Bonneville Estates Rockshelter In 2003 we began excavating an 8-m-long stratigraphic trench connecting the two block excavations. Through 2004, excavation of the trench penetrated to the top of the early-middle Holocene components discussed above. Besides well-preserved Fremontand Elko-aged cultural remains, our excavation of the trench has also yielded the first Clovis-aged artifact in the rockshelter. This is a Clovis fluted point base (38 mm wide and edge-ground) manufactured on green chert (Fig. 17D). It was found in the historic sheep dung deposit at the very top of the stratigraphic profile. Although undoubtedly redeposited, the primary context of the point may have been the rockshelter’s deeper sediments. Perhaps late prehistoric or even historic digging in Bonneville Estates Rockshelter led to its removal from that original place of deposition.

Figure 17. Lithic artifacts from Bonneville Estates Rockshelter. (A) Unhafted bifaces. (B–C and G) Stemmed point fragments. (D) Clovis fluted point fragment. (E–F and H) Unifacial scrapers.

Discussion Our excavations at Bonneville Estates Rockshelter will continue for several more years, and analyses of materials recovered from the excavations are in progress. Nonetheless, several preliminary conclusions can be made concerning early human activities in the rockshelter, based on initial observations of data recovered through 2004. Chronology The earliest unequivocal evidence for humans in Bonneville Estates Rockshelter is found in stratum 18b, dating to between 10,800 and 10,400 radiocarbon yr B.P. (Fig. 16). Possibly an earlier occupation is represented by stratum A19; however, additional excavations are needed to better define the age of this stratum as well as to unequivocally demonstrate whether lithic artifacts are primarily associated with faunal remains and features. Strata A18a and A17b in the West Block (Figs. 15 and 16) and strata 11/12 and 10 in the East Block (Fig. 18) are clearly related to the Great Basin’s stemmed point complex, but no diagnostic bifacial points have been encountered in stratum A18b. Nonetheless, the numerous hearths (n = 14) so far encountered in these strata indicate that Paleoindians frequently visited Bonneville Estates Rockshelter between ~10,800 and 9400 radiocarbon yr B.P. Shortly after 9400 radiocarbon yr B.P., however, these visits appear to have ceased, and Bonneville Estates Rockshelter was not occupied again until after 7500 radiocarbon yr B.P. The Danger Cave record has a similar hiatus in its 14C chronology (9800–8700 radiocarbon yr B.P.; see above), and along the Old

Figure 18. Stratigraphic profile of the East Block excavation of Bonneville Estates Rockshelter, exposed in 2001–2004.

River Bed, water appears to have ceased flowing by 8500 radiocarbon yr B.P. These data suggest that human population densities declined dramatically in the western Bonneville basin during the increasingly arid early Holocene. When humans returned to Bonneville Estates Rockshelter after 8000 radiocarbon yr B.P., they had adopted seeds into their diets, as shown by the presence of groundstone artifacts as well as pine-nut hulls and ricegrass seeds in the rockshelter’s early Archaic components.

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Technological Activities All of the pre-9400 radiocarbon yr B.P. debitage assemblages from Bonneville Estates Rockshelter are dominated by biface thinning flakes and retouch chips as well as complex platforms, indicating that the chief technological activities carried out in the rockshelter related to reduction and rejuvenation of tools, primarily bifaces. Nearly all debitage pieces are either tiny (<1 cm2) or small (1–3 cm2), further indicating that the bifaces being worked were far along in their reduction streams. Virtually none of the debitage pieces appear to relate to primary reduction activities, indicating that Bonneville Estates Rockshelter’s early inhabitants typically transported complete or nearly complete bifaces and unifaces to the rockshelter. In terms of lithic raw materials, Bonneville Estates Rockshelter’s early assemblages have about equal amounts of chert (41.7%) and fine-grained volcanics (e.g., basalt) (38.3%), while obsidian is less common (19.8%). Among 29 obsidian artifacts thus far subjected to X-ray fluorescence analysis, the majority originated from the Browns Bench (69.0%) and Malad (10.3%) sources, located 180 km north-northwest and 240 km northeast of Bonneville Estates Rockshelter, respectively. Other obsidians include Topaz Mountain, Utah, 105 km southeast (6.9%), Ferguson Wash, 6 km southeast (3.4%), and an unknown source (10.3%). These data indicate that the early inhabitants of Bonneville Estates Rockshelter were highly mobile, far-ranging foragers who made short but frequent stops at the rockshelter. Taphonomy and Subsistence Activities We have not recovered any remains of extinct fauna from the cultural components at Bonneville Estates Rockshelter. Nonetheless, faunal remains are abundant and well-preserved in Bonneville Estates Rockshelter’s lower cultural components. Taphonomic and zooarchaeological analyses of these materials are still incomplete, but some of the remains are clearly the product of human activity (e.g., sage grouse, artiodactyls, hare, bear), while others are more likely the product of nonhuman agents (e.g., cottontail, pygmy rabbit, rodent, waterfowl). The sage grouse remains are especially interesting. More than 250 grouse bones have been recovered that represent at least 20 sage grouse individuals. Four percent of the bones are charred and 12% display stone tool cut marks, leaving no doubt that the carcasses were systematically butchered and then discarded directly around the hearths of the early components. The artiodactyl bones identified belong primarily to pronghorn, but also include mountain sheep and deer. These bones were extensively broken and appear to be the result of marrow extraction. About one-third of these artiodactyl fragments were burned, and one displays stone tool cut marks. Regarding the hares, 24 tibiae, femora, and humeri cylinders have been identified, and 33% of these display stone tool cut marks. The faunal assemblages from Bonneville Estates Rockshelter are especially significant, in that they represent one of the first clear cases of butchered animal remains in association with archaeological artifacts pre-dating 10,000 radiocarbon yr B.P. in the Great Basin.

Paleobotanical analyses of hearth contents and other uncharred vegetal remains recovered from the early deposits of Bonneville Estates Rockshelter are still under way. However, preliminary results suggest that the rockshelter’s early human inhabitants did not regularly include seeds in their diets. Instead, virtually all of the plant macrofossils thus far identified in the pre-9000 radiocarbon yr B.P. components appear to have been the result of either humans collecting materials to burn in their fires, or nonhuman agents like woodrats. Plants do not appear to have become an important component of human diets at Bonneville Estates Rockshelter until the early Archaic, after 8000 radiocarbon yr B.P. SUMMARY The research discussed here, all conducted within the past five years, affords considerable new information about the nature and timing of earliest human occupation in the Bonneville basin, during an interval of substantial environmental change from the latest Pleistocene to the Holocene. In particular, this research demonstrates human occupation of the basin by at least 11,000 radiocarbon yr B.P., and perhaps by 12,300 radiocarbon yr B.P. The hundreds of occupation sites in the Old River Bed delta attest to an important subsistence focus in the basin’s wetland ecosystems, while faunal evidence from Bonneville Estates Rockshelter indicates a fairly broad array of prey species including large ungulates, hares, and sage grouse. Importantly, significant evidence for consumption of plant foods in these earliest occupation sites is presently lacking. The use of grinding stone technology for processing small seeds (a hallmark of the so-called Desert Archaic; Jennings, 1957) does not appear to have begun until after 8600 radiocarbon yr B.P., when the Old River Bed wetlands had dried up and the basin’s environment reached its modern desert character. Research planned for the next few years in the region will clarify the pattern of environmental change during the crucial latest Pleistocene-Holocene transition, confirm the dating of key occupations at Bonneville Estates Rockshelter and Old River Bed, and extend our analyses of the archaeological assemblages obtained from all of these sites. ACKNOWLEDGMENTS Our work at the Old River Bed has been supported by the U.S. Army, Dugway Proving Ground; we thank Kathleen Callister and Rachel Quist, Environmental Directorate, for their help and support. Our work at Danger Cave has been supported by the National Science Foundation (BCS-0312252), the Utah Division of State History, Utah Geological Survey, and the Lander Fund (Desert Research Institute). Our work at Bonneville Estates Rockshelter has been generously funded by the Sundance Archaeological Research Fund (University of Nevada, Reno), National Science Foundation, and U.S. Department of the Interior Bureau of Land Management.

Latest Pleistocene–early Holocene human occupation and paleoenvironmental change REFERENCES CITED Aikens, C.M., 1970, Hogup Cave: Salt Lake City, University of Utah Anthropological Papers 93, 286 p. Arkush, B.S., and Pitblado, B.L., 2000, Paleoarchaic surface assemblages in the Great Salt Lake desert, northwestern Utah: Journal of California and Great Basin Anthropology, v. 22, p. 12–42. Beck, C., and Jones, G.T., 1997, The terminal Pleistocene–Early Holocene archaeology of the Great Basin: Journal of World Prehistory, v. 11, p. 161–236. Beck, C., Taylor, A., Jones, G.T., Fadem, C.M., Cook, C.R., and Milward, S.A., 2002, Rocks are heavy: Transport costs and Paleoarchaic quarry behavior in the Great Basin: Journal of Anthropological Archaeology, v. 21, p. 481–507, doi: 10.1016/S0278-4165(02)00007-7. Beiswenger, J.M., 1991, Late Quaternary vegetational history of Grays Lake, Idaho: Ecological Monographs, v. 61, p. 165–182. Benson, L.V., Currey, D., Lao, Y., and Hostetler, S., 1992, Lake-size variations in the Lahontan and Bonneville basins between 13,000 and 9000 14 C yr B.P: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 95, p. 19–32, doi: 10.1016/0031-0182(92)90162-X. Bills, B.G., Wanbeam, T.J., and Currey, D.R., 2002, Geodynamics of Lake Bonneville, in Gwynn, J.W., ed., Great Salt Lake: An Overview of Change: Salt Lake City, Utah Department of Natural Resources Special Publication, p. 7–32. Bright, R.C., 1966, Pollen and seed stratigraphy of Swan Lake, southeastern Idaho: Its relation to regional vegetational history and to Lake Bonneville history: Tebiwa, v. 9, no. 2, p. 1–47. Broughton, J.M., 2000, The Homestead Cave ichthyofauna, in Madsen, D.B., Late Quaternary paleoecology in the Bonneville Basin: Salt Lake City, Utah Geological Survey Bulletin 130, p. 103–122. Broughton, J.M., Madsen, D.B., and Quade, J., 2000, Fish remains from Homestead Cave and lake levels of the past 13,000 years in the Bonneville basin: Quaternary Research, v. 53, p. 392–401, doi: 10.1006/qres.2000.2133. Carter, J.A., and Young, D.C., Jr., 2001, TS-5 central area and Craners cultural resource inventory, Wendover and Hill Air Force Ranges, Tooele and Box Elder Counties, Utah: Ogden, Utah, Hill Air Force Base, report on file, 243 p. Currey, D.R., 1980, Coastal geomorphology of Great Salt Lake and vicinity, in Gwynn, J.W., ed., The Great Salt Lake—a Scientific, Historical and Economic Overview: Salt Lake City, Utah Geological and Mineralogical Survey Bulletin 116, p. 69–82. Currey, D.R., 1982, Lake Bonneville: Selected features of relevance to neotectonic analysis: U.S. Geological Survey Open-File Report 82-1070, 31 p. Currey, D.R., 1990, Quaternary paleolakes in the evolution of semidesert basins, with special emphasis on Lake Bonneville and the Great Basin, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 76, p. 189– 214, doi: 10.1016/0031-0182(90)90113-L. Eardley, A.J., Gvosdetsky, V., and Marsell, R.E., 1957, Hydrology of Lake Bonneville and sediments and soils of its basin: Geological Society of America Bulletin, v. 68, p. 1141–1201. Elston, R.G., and Zeanah, D.W., 2002, Thinking outside the box: a new perspective on diet breadth and sexual division of labor in the Prearchaic Great Basin: World Archaeology, v. 34, p. 103–130, doi: 10.1080/ 00438240220134287. Fry, G.F., 1976, Analysis of prehistoric coprolites from Utah: Salt Lake City, University of Utah Anthropological Papers 97, 45 p. Godsey, H.S., Currey, D.R., and Chan, M.A., 2005, New evidence for an extended occupation of the Provo shoreline and implications for regional climate change, Pleistocene Lake Bonneville, Utah, USA: Quaternary Research, v. 63, p. 212–223. Goebel, T., Graf, K.E., Hockett, B.S., and Rhode, D., 2003, Late-Pleistocene humans at Bonneville Estates Rockshelter, eastern Nevada: Current Research in the Pleistocene, v. 20, p. 20–23. Graf, K.E., 2001, Paleoindian technological provisioning in the western Great Basin [M.S. thesis]: Las Vegas, University of Nevada, 197 p. Grayson, D.K., 1988, Danger Cave, Last Supper Cave, and Hanging Rock Shelter: The faunas: New York, American Museum of Natural History Anthropological Papers, v. 66, 130 p. Grayson, D.K., 1993, The desert’s past: A natural prehistory of the Great Basin: Washington, Smithsonian Institution Press, 356 p. Grayson, D.K., 1998, Moisture history and small mammal community richness during the latest Pleistocene and Holocene, northern Bonneville basin, Utah: Quaternary Research, v. 49, p. 330–334, doi: 10.1006/qres.1998.1970.

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Grayson, D.K., 2000, Mammalian responses to middle Holocene climatic change in the Great Basin of the western United States: Journal of Biogeography, v. 27, p. 181–192, doi: 10.1046/j.1365-2699.2000.00383.x. Haynes, C.V., 1991, Geoarchaeological and paleohydrological evidence for Clovis-age drought in North America and its bearing on extinction: Quaternary Research, v. 35, p. 438–450. Harper, K.T., and Alder, G.M., 1972, Paleoclimatic Inferences Concerning the last 10,000 years from a resampling of Danger Cave, Utah, in Fowler, D., ed., Great Basin cultural ecology: A Symposium: Reno, Nevada, Desert Research Institute Publications in the Social Sciences 8, p. 13–23. Huckleberry, G., Beck, C., Jones, G.T., Holmes, A., Cannon, M., Livingston, S., and Broughton, J.M., 2001, Terminal Pleistocene–early Holocene environmental change at the Sunshine locality, north-central Nevada, U.S.A.: Quaternary Research, v. 55, p. 303–312, doi: 10.1006/qres.2001.2217. Jarrett, R.C., and Malde, H.E., 1987, Paleodischarge of the late Pleistocene Bonneville flood, Snake River, Idaho: Geological Society of America Bulletin, v. 99, p. 127–134, doi: 10.1130/0016-7606(1987)99<127:POTLPB>2.0.CO;2. Jennings, J.D., 1957, Danger Cave: Salt Lake City, University of Utah Anthropological Papers 27, 328 p. Jennings, J.D., 1978, Prehistory of Utah and the eastern Great Basin: Salt Lake City, University of Utah Anthropological Paper 98, 263 p. Jones, G.T., Beck, C., Jones, E.E., and Hughes, R.E., 2003, Lithic source use and Paleoarchaic foraging territories in the Great Basin: American Antiquity, v. 68, p. 5–38. Madsen, D.B., 2000, Late Quaternary paleoecology in the Bonneville Basin: Salt Lake City, Utah Geological Survey Bulletin 130, 190 p. Madsen, D.B., 2002, Great Basin peoples and late Quaternary aquatic history, in Hershler, R., Currey, D.R., and Madsen, D.B., eds., Great Basin Aquatic Systems History: Washington, D.C., Smithsonian Contributions to Earth Sciences 33, p. 387–405. Madsen, D.B., and Rhode, D., 1990, Early Holocene piñon (Pinus monophylla) in the northeastern Great Basin: Quaternary Research, v. 33, p. 94–101, doi: 10.1016/0033-5894(90)90087-2. Madsen, D.B., Rhode, D., Grayson, D.K., Broughton, J.M., Livingston, S.D., Hunt, J.M., Quade, J., Schmitt, D.N., and Shaver, M.W., III, 2001, Late Quaternary environmental change in the Bonneville basin, western USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 167, p. 243–271, doi: 10.1016/S0031-0182(00)00240-6. Murchison, S.B., 1989, Fluctuation history of Great Salt Lake, Utah, during the last 13,000 years [Ph.D. dissertation]: Salt Lake City, University of Utah, 157 p. Murchison, S.B., and Mulvey, W.E., 2000, Late Pleistocene and Holocene shoreline stratigraphy on Antelope Island, Davis County, Utah, in King, J.K., and Willis, G.C., eds., Geology of Antelope Island: Salt Lake City, Utah Geological Survey Miscellaneous Publication 00-1, p. 77–83. Napton, L.K., 1997, The Spirit Cave mummy: Coprolite investigations: Nevada Historical Quarterly, v. 40, p. 97–104. O’Connor, J.E., 1993, Hydrology, hydraulics, and geomorphology of the Bonneville flood: Geological Society of America Special Paper 274, 90 p. Oviatt, C.G., 1988, Late Pleistocene and Holocene lake fluctuations in the Sevier Lake basin, Utah, USA: Journal of Paleolimnology, v. 1, p. 9–21, doi: 10.1007/BF00202190. Oviatt, C.G., 1997, Lake Bonneville fluctuations and global climate change: Geology, v. 25, p. 155–158, doi: 10.1130/0091-7613(1997)025<0155: LBFAGC>2.3.CO;2. Oviatt, C.G., Currey, D.R., and Sack, D., 1992, Radiocarbon chronology of Lake Bonneville, eastern Great Basin, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 99, p. 225–241, doi: 10.1016/00310182(92)90017-Y. Oviatt, C.G., Madsen, D.B., and Schmitt, D.N., 2003, Late Pleistocene and early Holocene rivers and wetlands in the Bonneville basin of western North America: Quaternary Research, v. 60, p. 200–210, doi: 10.1016/S00335894(03)00084-X. Oviatt, C.G., Miller, D.M., McGeehin, J.P., Zachary, C., and Mahan, S., 2005, The Younger Dryas phase of Great Salt Lake, Utah, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 219, no. 3-4, p. 263–284, doi: 10.1016/j.palaeo.2004.12.029. Quade, J., 2000, Strontium ratios and the origin of early Homestead Cave biota, in Madsen, D.B., Late Quaternary paleoecology in the Bonneville Basin: Salt Lake City, Utah Geological Survey Bulletin 130, p. 44–46. Rhode, D., 2000a, Middle and late Wisconsin vegetation in the Bonneville basin, in Madsen, D.B., Late Quaternary paleoecology in the Bonneville Basin: Salt Lake City, Utah Geological Survey Bulletin 130, p. 137–148.

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Rhode, D., 2000b, Holocene vegetation history in the Bonneville basin, in Madsen, D.B., Late Quaternary Paleoecology in the Bonneville Basin: Salt Lake City, Utah Geological Survey Bulletin 130, p. 149–164. Rhode, D., and Madsen, D.B., 1995, Early Holocene vegetation in the Bonneville basin: Quaternary Research, v. 44, p. 246–256, doi: 10.1006/qres.1995.1069. Rhode, D., and Madsen, D.B., 1998, Pine nut use in the early Holocene and beyond: The Danger Cave archaeobotanical record: Journal of Archaeological Science, v. 25, p. 1199–1210, doi: 10.1006/jasc.1998.0290. Rhode, D., Madsen, D.B., and Jones, K.T., 2006, Antiquity of early Holocene small-seed consumption and processing at Danger Cave, Utah, USA: Antiquity (in press). Sack, D., 1999, The composite nature of the Provo level of Lake Bonneville, Great Basin, western North America: Quaternary Research, v. 52, p. 316– 327, doi: 10.1006/qres.1999.2081. Schmitt, D.N., and Madsen, D.B., 2005, Camels Back Cave: University of Utah Anthropological Papers 125 (in press). Schmitt, D.N., Madsen, D.B., and Lupo, K.D., 2002, Small-mammal data on early and middle Holocene climates and biotic communities in the Bonneville basin, USA: Quaternary Research, v. 58, p. 255–260, doi: 10.1006/ qres.2002.2373.

Schroedl, A.R., and Coulam, N.J., 1989, Bonneville Estates Rockshelter: Elko, Nevada, Elko District Office, Bureau of Land Management, Cultural Resources Report 435-01-8906, 93 p. Steward, J.H., 1938, Basin-Plateau aboriginal sociopolitical groups: U.S. Bureau of American Ethnology Bulletin 120, 346 p. Thompson, R.S., 1990, Late Quaternary vegetation and climate in the Great Basin, in Betancourt, J.L., Van Devender, T.R., and Martin, P.S., eds., Packrat middens: The last 40,000 years of biotic change: Tucson, University of Arizona Press, p. 200–239. Thompson, R.S., 1992, Late Quaternary environments in Ruby Valley, Nevada: Quaternary Research, v. 37, p. 1–15. Thompson, R.S., Whitlock, C., Bartlein, P.J., Harrison, S.P., and Spaulding, W.G., 1993, Climate changes in the western United States since 18,000 yr B.P., in Wright, Jr., H.E., et al., eds., Global climates since the Last Glacial Maximum: Minneapolis, University of Minnesota Press, p. 468–513. Willig, J.A., Aikens, C.M., and Fagan, J.L., editors, 1988, Early human occupation in far western North America: The Clovis-Archaic interface: Carson City, Nevada State Museum Anthropological Papers 21, 482 p.

Printed in the USA

Geological Society of America Field Guide 6 2005

Neotectonics and paleoseismology of the Wasatch fault, Utah Ronald L. Bruhn* Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112, USA Christopher B. DuRoss* Utah Geological Survey, 1594 West North Temple, #3110, Salt Lake City, Utah 84116, USA Ronald A. Harris* Department of Geology, Brigham Young University, Provo, Utah 84602, USA William R. Lund* Utah Geological Survey, Southern Utah Regional Office, 88 East Fiddler Canyon Road, Cedar City, Utah 84720, USA

ABSTRACT The Wasatch fault is a 370-km-long zone of normal faulting that forms the eastern edge of the Basin and Range Province in Utah and southeastern Idaho. The fault zone is subdivided into ten segments that range from 30 to 60 km in length and are each capable of generating earthquakes of M ~7. For the five central segments, multiple surface-faulting earthquakes have occurred during the Holocene, and vertical slip rates are ~1 mm/yr. Recurrence intervals for the individual central segments range from ~1300 to 2500 yr. The fault poses a significant seismic hazard to the highly urbanized Wasatch Front in northcentral Utah. The field localities described in this guide provide an overview of the surface and subsurface character of the Wasatch fault zone. Five field trip stops are located along the Nephi, Provo, and Salt Lake City segments. We will observe fault scarps on Quaternary deposits, which record tectonic displacements associated with Holocene earthquakes, and fault-zone rocks exhumed from depths in excess of 10 km that are hydrothermally altered and have evidence of brittle and ductile deformation. Specific topics of discussion include the nature of piedmont fault scarps; the use of paleoseismic trenching and faultscarp geomorphology to infer earthquake timing, recurrence intervals, and fault slip rates; and the subsurface structure and rheology of the fault. Keywords: faulting, paleoseismology, neotectonics, earthquake geology. INTRODUCTION The Wasatch normal fault is one of the preeminent tectonic structures of North America. The fault has a profound physio*E-mails [email protected]; [email protected]; ron_harris@ byu.edu; [email protected].

graphic expression on the landscape of the western United States and poses a significant earthquake hazard to the densely populated urban corridor of north-central Utah. The Wasatch fault played a central role in the evolution of geological thought in the latter nineteenth and early twentieth centuries, when geologists began pondering the landscape and geological structure of the Great Basin (Emmons, 1878; King, 1878; Gilbert, 1884, 1928;

Bruhn, R.L., DuRoss, C.B., Harris, R.A., and Lund, W.R., 2005, Neotectonics and paleoseismology of the Wasatch fault, Utah, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 231–250, doi: 10.1130/2005.fld006(11). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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Pack, 1926). Subsequently, the Wasatch fault regained prominence in the latter part of the twentieth century when it played a significant role in the development of modern paleoseismology. Studies of the Wasatch fault contributed directly to the concept of “characteristic” earthquakes (Schwartz and Coppersmith, 1984) and to our understanding of colluvial-wedge formation, the recurrence patterns of normal faulting (e.g., Machette et al., 1992; McCalpin, 1996; Lund, 2005), and the mechanics of earthquake generation (Parry and Bruhn, 1986; Bruhn et al., 1990, 1994). Today, the fault remains an important natural laboratory for investigating the structure, tectonic geomorphology, earthquake-recurrence patterns, and rupture segmentation of active normal faults. Considerable effort is currently directed at reconciling differences in slip rates inferred from Quaternary geology versus geodetic measurements, a topic of fundamental importance to tectonophysics as well as the field of earthquake hazards (Chang and Smith, 2002; Friedrich et al., 2003). The topographic escarpment of the Wasatch Range rises abruptly along the eastern side of the Wasatch fault (Fig. 1), forming the physiographic boundary between the Rocky Mountains or the Colorado Plateau and the Basin and Range Province to the west. Evidence for active faulting is both abundant and well displayed where piedmont fault scarps cross Quaternary alluvial fans (Fig. 2), glacial moraines, and deltas and shorelines of Pleistocene Lake Bonneville. Paleoseismic trenching of these faulted deposits reveals a rich and varied history of late Pleistocene and Holocene earthquakes, providing a tantalizing glimpse into temporal patterns of earthquake recurrence—the “pulse” of the earthquake engine. The geomorphology and structure of the Wasatch mountain front also contain clues as to the geometry, composition, and mechanical characteristics of the fault at depth. When taken collectively, this information allows us to peer into the heart of the earthquake engine itself. During our field trip, we will visit localities along three of the ten proposed rupture segments of the Wasatch fault (Fig. 1). The goal is to observe and discuss the fault from the surface to depths in excess of 10 km, where M 7 earthquakes presumably originate. We will observe and discuss aspects of the tectonic geomorphology and Quaternary geology, as well as visit a paleoseismic trench excavation and localities where deformed rocks are exhumed from deep within the fault zone. Specific topics considered include the nature of piedmont fault scarps, the use of paleoseismic trenching and geomorphology surveys to infer fault-rupture parameters (e.g., earthquake timing, recurrence, and displacement), and the structure and rheology of the fault in the subsurface. WASATCH FAULT—TECTONICS AND STRUCTURE The Wasatch normal fault extends for ~370 km from southern Idaho into central Utah, passing directly through the densely populated and rapidly expanding urban corridor along the front of the Wasatch Range. The fault is subdivided into ten segments that vary from roughly 30–60 km in length and that

were originally thought to be independent rupture segments based on differences in Holocene earthquake histories (Schwartz and Coppersmith, 1984). This concept is currently being challenged as more information on earthquake timing is obtained and the implications of event-dating uncertainties are further explored (Chang and Smith, 2002; DuRoss and Bruhn, 2005; Lund, 2005). The segment boundaries are also interesting from a structural perspective because they are marked by bedrock spurs that extend into adjacent basins, by changes in the topography along the Wasatch Range crest, and by jogs and bends in the fault trace (Gilbert, 1928; Schwartz and Coppersmith, 1984; Bruhn et al., 1992; Wheeler and Krystinik, 1992). Each segment is capable of generating earthquakes of M ~7. Although no such earthquakes have occurred in historical time, the six central segments (Brigham City to Levan segments; Fig. 1) have ruptured in large magnitude earthquakes that created surface displacement during the Holocene. Recurrence intervals of surface-faulting earthquakes vary from segment to segment, but for the five central segments (not including the Levan segment) average roughly 1300 yr to 2500 yr in each segment, or about once every three to five hundred years when considered collectively (Machette et al., 1992; McCalpin and Nishenko, 1996; Lund, 2005). Surface displacement or net vertical tectonic displacement (Swan et al., 1980) averages ~2 m during surface-faulting earthquakes, but varies along the length of the rupture. Structurally, the Wasatch fault is an aggregate of normal faults that nucleated, grew in length, and linked together to form a continuous normal fault zone. The process of growth and linkage is reflected in the geometry of the ten fault segments (Wheeler and Krystinik, 1992), variations in the amount of cumulative displacement along strike, differences in the amount of flexure and rotation of the footwall, and the structure of hanging wall basins (e.g., Zoback, 1983; Mabey, 1992; Armstrong et al., 2003, 2004). The initiation of faulting is partly constrained by radiometric dating of fault rock (Parry et al., 1988), by the onset of Neogene sedimentation in the eastern Great Basin (Bryant et al., 1989), and by the thermochronology of the Wasatch Range (e.g., Evans et al., 1985; Kowallis et al., 1990). The central part of the Wasatch fault may have formed as early as 17.6 ± 0.7 Ma based on a K-Ar age of sericite obtained from the fault zone near Salt Lake City, where deformed rocks were exhumed from a depth of ~11 km (Parry and Bruhn, 1987). By 10–12 Ma, faulting was clearly under way, based on an influx of sedimentary detritus into basins west of the Wasatch Range and the fission track thermochronology of the Wasatch Range (Armstrong et al., 2003, 2004). Vertical slip rates averaged over the past 10 Ma vary from 0.4 to 0.7 mm/yr, somewhat lower than the 1 mm/yr average obtained from fault trench studies (Machette et al., 1992; Friedrich et al., 2003; Lund, 2005). The contemporary rate of horizontal extension measured by geodetic surveying (using GPS networks) is 2.7 ± 1.3 mm/yr in the central part of the fault zone (Martinez et al., 1998). If accounted for by seismic processes, the Wasatch fault geodetic rate would suggest a higher frequency of large (M ≥ 6.6) earthquakes than supported by paleoseismic and historical-seismicity

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Figure 1. Shaded-relief map of the Wasatch Front with trace of the Wasatch fault (from Black et al., 2003), the Salt Lake City, Provo, and Nephi fault segments, and the locations of paleoseismology trenches. SP—Salt Palace; LC—Little Cottonwood Canyon; SF—South Fork Dry Creek; DC—Dry Gulch; AF—American Fork Canyon; RCr—Rock Creek; HC—Hobble Creek; MP—Mapleton; WC—Water Canyon; WH—Woodland Hills; NC— North Creek; and RCyn—Red Canyon). Field trip starting location (S) and stops 1 through 5 are labeled and marked by black dots. Inset map shows Wasatch fault zone (WFZ) and six central segments: BC—Brigham City; WB—Weber; SLC—Salt Lake City; PV—Provo; NP—Nephi; LV—Levan.

observations, possibly indicating the accommodation of strain by aseismic processes (Chang and Smith, 2002). Direct comparisons between horizontal rates of extension determined from geodetic data and vertical tectonic rates estimated by geologists requires knowledge of the fault’s subsurface dip, which is a challenging and unresolved problem.

The subsurface geometry of the Wasatch fault has become a topic of discussion and controversy ever since G.K. Gilbert (1928) concluded that the fault dipped between 30° and 40° with an average dip of ~33°. Gilbert (1928) based his conclusion on the geometry of exposed fault surfaces and breccia zones exposed along faceted mountain fronts. He emphasized that the more

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R.L. Bruhn et al. wall, they form narrow east-west–trending belts that extend westward along segment boundaries. FIELD TRIP Stop 1: Red Canyon Fault Scarp—Nephi Segment Travel Directions to Stop 1 Depart from the Salt Palace Convention Center and drive south on West Temple to 500 South. Turn right (W) on 500 South and enter the I-15 freeway driving south toward Provo and Las Vegas. Proceed south for 81 mi to Exit 228 for Nephi, Utah. After exiting, turn left (E) and proceed over the freeway and then right (S) onto the frontage road that parallels the freeway. Proceed south for 0.8 mi and park in front of the corral.

Figure 2. The Red Canyon fault scarp locality near the southern end of the Nephi segment of the Wasatch fault. The view is toward the east with I-15 in the middle ground. Note the increase in vertical displacement (VD) across the scarp from ~6 m on the Holocene alluvial surface (H) to ~20 m across the late Pleistocene fan remnant (P). RCyn—Red Canyon fault trench site (Jackson, 1991).

steeply dipping faults beneath piedmont scarps were probably high-angle splays that emanated and propagated upward from the main fault surface at shallow depths. Subsequent work has shown that Gilbert’s (1928) observations remain valid. However, few subsurface data have been obtained in the intervening years despite the advent of reflection seismology, and few drill holes have penetrated the fault at depth. An exploration well penetrated the fault near Brigham City where the fault dips ~45° west (W.A. Yonkee, 2005, personal commun.). Published reflection-seismic data across the Wasatch fault are 1970s in vintage that are difficult to interpret and mostly provide evidence for the dips of Tertiary basin sediments (Smith and Bruhn, 1984). Recent geodetic measurements suggest creep on a low-angle fault or shear zone at depths of 10 km to 15 km based on numerical modeling (W.L. Chang, 2005, personal commun.). The fault-dip angle at shallower levels is better constrained, based on the attitude of exposed fault surfaces, which vary from moderate (30–40°) to more steeply dipping at some localities (Harris et al., 2000). Historical seismicity surrounding the Wasatch fault is of particular interest both for exploring the structure and geodynamics of the fault and for detecting changes in seismicity that may precede a large earthquake. The seismicity reveals little about the subsurface structure of the fault. Earthquakes are small- to moderate-magnitude events that are scattered over a broad area. In fact, most earthquakes are located beneath the Wasatch Range to the east of the fault, rather than beneath basins in the hanging wall (Arabasz and Julander, 1986; Arabasz et al., 1979, 1992). Where concentrations of earthquakes do occur in the hanging

Background The Nephi segment consists of fault scarps extending 42 km from Nephi to north of Payson along two distinct strands: the 17-km-long Santaquin strand and 25-km-long Nephi strand (Machette et al., 1992; DuRoss and Bruhn, 2005; Fig. 3). The northern Santaquin strand bounds the west side of Dry Mountain, and the southern Nephi strand bounds the Wasatch Range east of Juab Valley. A ~6-km-long, northeast-striking cross fault connects the Nephi and Santaquin strands (Machette et al., 1992; Harty et al., 1997; DuRoss, 2004). DuRoss (2004) found no evidence of Holocene or late Pleistocene displacement along the cross fault, but heavy vegetation and landslides may obscure late Quaternary displacement. The Nephi segment is separated from the Provo segment to the north by a 4.5–9-km-wide en echelon right step, and from the Levan segment to the south by a 15-km gap in Holocene and latest Pleistocene surface faulting (Hylland and Machette, 2004). The stop at Red Canyon provides an opportunity to view a multiple-event fault scarp that increases in height where it crosses from Holocene to late Pleistocene alluvial fan surfaces (Fig. 2). Fault slip rates estimated from trenching of the Holocene fan and scarp-diffusion modeling of the fault scarps on a late Pleistocene fan remnant indicate that earthquake recurrence may have been temporally clustered, with a higher rate of activity during the Holocene (Machette et al., 1992; Mattson and Bruhn, 2001; DuRoss and Bruhn, 2005). Fault-Scarp Geomorphology and Slip Rate Nonlinear scarp-diffusion modeling simulates the diffusive erosion (e.g., by rain splash and soil creep) of the upper half of a measured scarp profile to determine the time at which surface displacement began, or the scarp initiation time (Andrews and Bucknam, 1987; Mattson and Bruhn, 2001; DuRoss and Bruhn, 2005). The diffusion models used for the Nephi-segment scarp profiles used a diffusivity constant calibrated with Bonneville shoreline and fault-trench data (Mattson and Bruhn, 2001), and accounted for geologic and scarp-profile instrumentation uncertainty by running numerous simulations with randomized

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Figure 3. Nephi segment of the Wasatch fault, showing the Nephi and Santaquin fault strands and locations of previous paleoseismology trenches (annotations as per Fig. 1) and 2005 trench sites (white triangles: SQ—Santaquin site; WCr—Willow Creek site). Inset figure shows location of Stop 2. Solid black line indicates Wasatch fault trace; ball and bar on downthrown side (Machette, 1992; Harty et al., 1997; Black et al., 2003).

model parameters for each profile (DuRoss and Bruhn, 2005). Following the simulations, DuRoss and Bruhn (2005) calculated the mean vertical displacement and best-fit scarp initiation times, and formulated multiple rupture scenarios for the Nephi segment by integrating fault-trace and existing paleoseismic information. DuRoss and Bruhn (2005) calculated slip rates based on the scarpdisplacement data and rupture scenarios, dividing the vertical displacement associated with a single earthquake by the elapsed time since the previous event, or in the case of multiple earthquakes, the average vertical displacement per event divided by the average recurrence interval between events. Slip rates calculated by Mattson and Bruhn (2001) include the elapsed time since the most recent earthquake and are referred to as open-ended rates.

Fault-scarp profiles and trench data (Jackson, 1991) indicate that vertical displacement across the fault scarp near Red Canyon increases from ~6 m on a Holocene alluvial fan surface to 20 m on the late Pleistocene fan remnant (Fig. 2). A fault trench excavated across the 6-m scarp revealed evidence for three earthquakes in the past 5–6 ka (Jackson, 1991) and a vertical slip rate of 0.6–1.0 mm/yr, based on scarp displacement, thickness of colluvial-wedge deposits, and earthquake timing (Harty et al., 1997; Lund, 2005). Surface displacement associated with these Holocene earthquakes is localized along the base of the 20-m scarp farther north, where there is an abrupt increase in slope angle along the lower part of the scarp’s topographic profile (Mattson and Bruhn, 2001). The 20-m scarp may represent at least nine

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surface-faulting earthquakes, assuming an average 1.8 ± 0.4 m of vertical displacement per event, based on the preferred faultrupture scenario of DuRoss and Bruhn (2005). Topographic profiling and scarp-diffusion modeling of the 20-m and 6-m scarps indicate that displacement across the late Pleistocene alluvial fan surface initiated at ca. 70 ka, yielding an open-ended slip rate of 0.4 mm/yr, less than one-half of the Holocene open-ended slip rate (1.3 mm/yr since 4.3 ka, based on diffusion modeling), possibly due to earthquake clustering during the Holocene (Mattson and Bruhn, 2001). DuRoss and Bruhn (2005) modeled the initiation time of similar scarps elsewhere along the Nephi segment, and calculated a late Pleistocene slip rate of 0.2 mm/yr (based on scarp displacement that occurred between 10–15 ka and 48–59 ka), in contrast to a Holocene slip rate of 0.5–0.7 mm/yr. The variations in slip rate determined at Red Canyon and along the Nephi segment depend on the scarp-modeling assumption that surface displacement accumulates along a single fault beneath the middle of the fault scarp (Mattson and Bruhn, 2001). This assumption needs to be reconsidered if preliminary results of our recent seismic-imaging experiments are valid. The seismic reflection image shows three faults displacing Quaternary deposits near the top, middle, and base of the 20-m scarp (Fig. 4). Subsurface imaging does not yet allow us to discriminate between rupture scenarios with distributed faulting on two or more of the buried faults during single earthquakes versus the possibility that

Figure 4. Seismic reflection image of the 20-m scarp at Red Canyon. Three faults are inferred beneath the top, middle, and base of the scarp (white arrows). The seismic section crosses the scarp near the ~20 m vertical displacement annotation in Figure 2. Data collection and processing provided by T. Crosby, University of Utah, using 72-fold CMP reflection processing with statics corrections for elevation and anomalous near-surface velocity.

fault activity migrated toward the base of the scarp with time. These potential scenarios are critical to properly interpreting the evolution of the scarp profile when solving for slip rate and the time of scarp initiation on the late Pleistocene fan. Stop 2: Nephi Segment Fault Trench Travel Directions to Stop 2 Return to I-15 and proceed north to the Santaquin exit (244). Turn right (E) at the end of the exit ramp and continue north on the frontage road for 0.7 mi. Turn right (NE) into a residential subdivision and continue 0.1 mi to a small city park. Park here and walk ~200 m SE along a small stream channel to the Wasatch fault (Fig. 3). Introduction In May of 2005, the Utah Geological Survey (UGS) and U.S. Geological Survey (USGS) excavated trenches at two sites along the Nephi segment to develop information on the timing and recurrence of earthquakes and segmentation of surface faulting. The sites include the Santaquin site (UGS), ~2 km east of Santaquin, and the Willow Creek site (USGS), ~8 km north of Nephi (Fig. 3). At Stop 2 we will visit fault trenches at the Santaquin site, located along the northern part of the Nephi segment near the mouth of Santaquin Canyon. At the site, prehistoric surface faulting has vertically displaced a late Holocene alluvial fan surface 2.5 m down-to-the-west, making it an ideal location to investigate the most recent surface faulting on the northern part of the Nephi segment. Because this field trip guide was written before excavation of the Santaquin trenches, we herein provide background information and discuss the reasons for the trench investigations. Background The Nephi segment has an extensive record of latest Quaternary surface faulting along its northern and southern fault strands but arguably the most poorly constrained earthquake chronology of the central, active segments of the Wasatch fault (Lund, 2005). Fault-scarp diffusion analyses along the Nephi segment indicate that at least six large-magnitude (average 6.5–7.1) earthquakes have ruptured all or part of the segment since ca. 12 ka (DuRoss and Bruhn, 2005). Most multiple-event scarps along the segment have <10 m of vertical displacement, although scarps on the oldest (~late Pleistocene) alluvial fans and landslides indicate a long history of earthquake activity with 11–25 m of displacement over the past several tens of thousands of years (DuRoss, 2004). Despite the long record of earthquakes preserved along the fault trace, trench studies (Hanson et al., 1981; Jackson, 1991) have identified only three poorly constrained mid- to late Holocene earthquakes on the southern part of the segment, supporting the need for additional fault trench investigations. The motivation for new trenching along the northern and southern parts of the Nephi segment is that understanding the fault rupture parameters (e.g., timing, displacement, and extent)

Neotectonics and paleoseismology of the Wasatch fault for the individual fault strands is critical to resolve the overall earthquake behavior of the Nephi segment. In addition, paleoseismic data for the segment are presently (April 2005) limited to the Nephi strand and poorly constrain three surface-faulting earthquakes. Regional scarp analyses suggest a potential for both partial and multi-segment ruptures along the Nephi segment (DuRoss and Bruhn, 2005), but seismic-hazard analyses assume the two strands form a single segment, and as a result rely on the poorly constrained earthquake record from Nephi strand fault trenches. Thus, the UGS and USGS excavated trenches near Santaquin and Willow Creek to (1) clarify the history of paleoearthquakes on the Nephi strand (Willow Creek site), (2) determine the timing of paleoearthquakes on the Santaquin strand (Santaquin site), (3) accurately characterize the seismic-source potential of the Nephi segment, and (4) compare the Nephi segment paleoearthquake parameters (timing, displacement, recurrence) with those for the adjacent Provo and Levan segments. Information on earthquake timing obtained at the Santaquin trench site should resolve both the nature of the Nephi-Provo segment boundary and the relation of the Santaquin strand to the Nephi strand, when compared with new paleoseismic data developed at Willow Creek. Resolving the individual rupture histories of the two strands is important, as asynchronous ruptures would result in more frequent moderate to large magnitude earthquakes, whereas synchronous ruptures of the strands would generate less frequent but larger earthquakes. Also, simultaneous ruptures of the Provo segment and Santaquin strand of the Nephi segment would create larger magnitude earthquakes than previously expected. Ultimately, the new trench information will serve to refine fault segmentation and seismic hazard models for the Wasatch fault, and improve our understanding of fault segments as individual and combined seismic sources. Earthquake Chronology As part of a review of paleoseismic data for Quaternary faults in Utah, Lund (2005) reported a consensus earthquake chronology for the Nephi segment (Table 1) and described the earthquake-timing information on the Nephi segment as the most poorly constrained of the five active central segments of the Wasatch fault. This Nephi segment chronology is based on fault trench studies at North Creek (Hanson et al., 1981) and Red Canyon (Jackson, 1991) on the Nephi strand, which identified evidence for three surface-faulting earthquakes since the middle Holocene (based on fan alluvium 14C dated at ca. 5300 cal yr B.P. by Bucknam, 1978). However, the investigations produced conflicting sets of numerical ages, and as a result, significantly different earthquake chronologies exist depending on how the age data are treated. For example, at North Creek, five 14C ages for a soil formed on top of the event Y colluvial wedge cluster into two groups between 1400 and 1650 14C yr B.P. and between 3600 and 3900 14C yr B.P. Hanson et al. (1981) preferred the older sets of ages as representative of the soil and the minimum limit on the timing of event Y. At the Red Canyon site over 12 km south of the

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North Creek site, Jackson (1991) selected 14C and thermoluminescence ages for the two most recent events that were consistent with the North Creek data. He extrapolated the age of the North Creek fan alluvium (Bucknam, 1978) in order to constrain the timing of event X at Red Canyon. DuRoss and Bruhn (2005) formulated multiple earthquake scenarios for the Nephi segment based on scarp diffusion modeling, scarp trace, and vertical displacement data (Table 2; Fig. 5). The preferred earthquake history includes six ruptures on the Nephi segment since the latest Pleistocene, having an average vertical displacement per event of 1.8 m. Two events ruptured the entire Nephi segment (both Nephi and Santaquin strands simultaneously), and four younger events ruptured only one of the fault strands. The earthquake time ranges are qualitative estimates, based on diffusion model and fault-trench data uncertainties. The broad ranges and possible partial-segment ruptures illustrate the need for additional paleoseismic data to test and constrain the timing of events along both fault strands. Earthquake Rupture Patterns Partial segments. Based on the earthquake timing (from both fault-trench and diffusion modeling data), along-strike changes in scarp morphology and displacement, and surface

TABLE 1. CONSENSUS EARTHQUAKE TIMING FOR THE NEPHI SEGMENT Event

Consensus time range

Z

<1 ± 0.4 ka, but possibly as young as 0.4 ± 0.1 ka

Y

ca. 3.9 ± 0.5 ka

X

>3.9 ± 0.5 ka, <5.3 ± 0.7 ka

Note: From Lund (2005).

TABLE 2. PREFERRED EARTHQUAKE CHRONOLOGY FOR THE NEPHI SEGMENT BASED ON SCARP-DIFFUSION MODELING AND FAULT-TRENCH DATA. Event*

Preferred time range

ZS

0.4–0.6 ka (partial rupture—Santaquin strand?)

ZN

1.0–1.4 ka (partial rupture—Nephi strand?)

YS

2.0–3.0 ka (partial rupture—Santaquin strand?)

N

Y

ca. 3.9–4.0 ka (partial rupture—Nephi strand?)

X

ca. 5.2–7.0 ka

W

ca. 10-15 ka

Note: Modified from DuRoss and Bruhn (2005). *Events ZS and YS are based on diffusion-modeling data. Events ZN, YN, and X are based on both diffusion-modeling and fault-trench data. Event YN is based primarily on paleoseismic-trench data due to a paucity of scarps recording only the youngest two events on the Nephi strand. Event W is inferred from displacement across Bonneville shoreline.

R.L. Bruhn et al. the Provo segment to the north (DuRoss and Bruhn, 2005). Provo segment ruptures propagating toward the fault’s southern tip would need to transfer slip across a 4.5–6.5-km distance to rupture the Santaquin strand in Santaquin Canyon or a ~9-km distance to rupture the northern part of the Santaquin strand (Fig. 3). DuRoss and Bruhn (2005) reported that the youngest ruptures on the Santaquin strand are mostly limited to Santaquin Canyon (~7 km long; Figs. 3 and 5) and scarps associated with the events initiated at ca. 0.5 ka and ca. 2.6 ka based on diffusion modeling. In comparison, fault trench studies on the Provo-segment identified large displacement ruptures along the entire fault, the youngest of which occurred ca. 0.6 and ca. 2.85 ka (Lund et al., 1991; Lund and Black, 1998; Lund, 2005). Based on mechanical modeling, a 60-km-long rupture of the Provo segment would induce the greatest Coulomb failure stress change on the southern half of the Santaquin strand (four bars at 10 km depth; Chang, 1998), decreasing significantly to the south along the Nephi strand. Conversely, DuRoss and Bruhn (2005) considered the Payson salient to be an effective barrier to ruptures propagated on the Nephi segment, based on the fault geometry, distribution of slip, and minor stress change

fault–trace geometry, DuRoss and Bruhn (2005) suggested that the 4.2- to 7-km-wide Nephi-Santaquin fault-strand step-over may impede small ruptures on the individual strands, evidenced by the proposed timing and geometry of rupturing along the two strands (DuRoss and Bruhn, 2005). Along both strands, the youngest fault rupture traces bifurcate and the vertical displacements decrease abruptly toward the strand step-over, indicating rupture impediment. The youngest two ruptures on each fault strand are morphologically and structurally distinct across the strand step-over. The youngest displacement on the Santaquin strand is limited to the southern half of the strand where 1.0 m of displacement occurred at ca. 0.4–0.6 ka, based on scarp diffusion modeling (DuRoss and Bruhn, 2005). In contrast, the youngest rupture on the Nephi strand occurred at ca. 1.0–1.4 ka with 1.7 m of displacement, based on diffusion modeling and fault trench data. Scarp initiation times suggest that the next oldest event on the Santaquin strand occurred at ca. 2.0–3.0 ka, in contrast to the event on the Nephi strand estimated at ca. 3.9–4.0 ka, based dominantly on fault trench data. Multiple segments. The Santaquin strand may also rupture due to coseismic stress interaction with paleoearthquakes along

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CUMULATIVE VERTICAL DISPLACEMENT

Vertical displacement (m)

25

20

Alluvial fan MRE: Af1: Af2y: Af2o: Af3:

Santaquin strand 1.0 ± 0.1 m 3.0 ± 0.8 m 5.3 ± 1.1 m 8.1 ± 1.9 m 14.7 ± 2.0 m

Nephi strand 1.7 ± 0.5 m 3.7 ± 0.4 m 5.5 ± 1.1 m 7.2 ± 0.7 m 16.6 ± 2.2 m

Mean scarp initiation time from scarps on Af3 fan surfaces along Nephi strand Af3

?

15

?

?

?

52.4 ± 4.9 ka

?

Af3 54.0 ±8.1 ka

?

10 2.6 ±0.7 ka 0.5 ±0.1 ka

5

0

5

MF

?

? Af2y

?

7.0 ±1.4 ka

?

4.0 ±1.5 ka

15

Af2y Af1

?

1.4 ±0.5 ka MRE

MRE

10

Af2o

?

? Af1

Af2o

0

11.4 ±1.9 ka

13.8 ± 3.7 ka

20

25

30

35

40

Distance along fault (km from N fault tip) Figure 5. Distribution of vertical displacement along the Nephi segment from displaced alluvial fan surfaces. Diamonds—MRE (most recent event); Af—alluvial fan: triangles—Af1; squares—Af2 (y—younger, o—older); stars—Af3 (modified from DuRoss and Bruhn, 2005). Gray shapes and inset chart of vertical displacement data from scarp-profile surveying (DuRoss and Bruhn, 2005); black shapes from Hanson et al. (1981), Jackson (1991), Machette (1992), and Mattson and Bruhn (2001). Mean scarp initiation times (±1σ) from scarp-diffusion modeling (DuRoss and Bruhn, 2005). MF—Mendenhall fault.

Neotectonics and paleoseismology of the Wasatch fault (~1.8 bars; Chang, 1998) imposed on the southern Provo segment due to a Nephi segment earthquake. Stop 3: Provo Segment of the Wasatch Fault and the Seven Peaks Fault Scarp Travel Directions to Stop 3 Return to I-15 and drive north 18 mi to Exit 263 (Utah Hwy 189). Proceed north on University Avenue for 1.5 mi and then turn right onto 300 South. Continue east on 300 South for 0.8 mi to the intersection with 900 E. Turn left at 900 E and drive 0.3 mi north to East Center Street. Turn right onto East Center St. and drive 0.1 mi to Seven Peaks Blvd. Proceed northward for 0.3 mi to the Seven Peaks Resort. An exposure of the Wasatch fault is located at the base of the mountain just east of the resort headquarters. Provo Segment of the Wasatch Fault The Provo segment of the Wasatch fault bounds the eastern side of Utah Valley (Figs. 1 and 6). It has a 60-km-long surface rupture trace that extends from the Traverse Mountains south

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of Salt Lake City to Payson. The most recent major earthquake (moment magnitude [MW] ~7.1) along the Wasatch fault occurred along the Provo segment between 500 and 700 yr B.P., based on fault trenching (Lund et al., 1991; Machette, 1992; Lund and Black, 1998). Scarps from this earthquake are still well preserved in places. Displacement across the Bonneville shoreline (16.8–18 ka, based on radiocarbon dating; D. Currey, written commun. in Lund, 2005) varies along the segment from 15 to 26 m near American Fork (Machette, 1992) to 40–45 m near Hobble Creek (Swan et al., 1980; Machette et al., 1992; Lund, 2005), which yields long-term (open-ended) slip-rate estimates of 0.8–1.5 mm/yr and 2.2–2.7 mm/yr, respectively (Machette, 1992; Lund, 2005). Contemporary horizontal loading rates for the fault vary from 2 to 4 mm/yr, based on geodetic measurements (EDM and GPS) over the past few decades (Martinez et al., 1998; Friedrich et al., 2003). Differences between these rates can be accounted for either by movement along a low-angle slip surface or by viscoelastic behavior, in which a significant amount of the viscous component is not recovered during coseismic deformation (Harris et al., 2000).

Figure 6. Block diagram of the Wasatch Front and Utah Valley near Rock Canyon (E-W line) and Lindon (N-S) line. The Wasatch fault cuts through the back limb of the Rock Canyon anticline, which plunges north. The ramp associated with this fault propagation–type fold may have been reactivated as part of the Wasatch fault. A series of antithetic faults with top-down-to-the-east slip offset the folded rocks in the footwall. Cross section constructed by R. Harris and E. Robeck.

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The footwall of the Provo segment consists of mostly Paleozoic and locally Precambrian sedimentary rocks (Fig. 6). In most cases these rocks form part of the gently dipping back limb of the east-verging Nebo-Charleston thrust sheet. Structural reconstructions indicate that the Wasatch fault is parallel to, and may have reactivated thrust ramps associated with, the Nebo-Charleston thrust (Smith and Bruhn, 1984). Many of the gently dipping sedimentary rocks in the footwall are tilted up to 40° west by drag along the Wasatch fault (Fig. 7). The break in slope on the footwall evident throughout the Provo segment marks the position of the erodable Upper Mississippian to Lower Pennsylvanian Manning Canyon Shale, in contrast to the cliffs of overlying Pennsylvanian to Permian Oquirrh Group limestone and sandstone (Fig. 6). The hanging wall of the Provo segment consists mostly of blocks of Paleozoic rocks and of upper Cenozoic basin fill. Most of the blocks consist of Mississippian limestone, in some cases displaced thousands of meters downward from their footwall source (Fig. 6). Middle Miocene to modern basin-fill deposits are estimated to be a maximum of 4 km thick from drilling and gravity surveys (Mabey, 1992). The stratigraphic thickness of hanging-wall rocks below the basin fill and above the Wasatch fault is estimated at ~10 km (Hintze, 1988). This implies that there has been at least 10 km of vertical displacement along the fault since the middle Miocene. The bedrock part of the hanging wall emerges from the basin fill to the west of Utah Valley to form Lake Mountain. This configuration is consistent with hanging wall rotation associated with listric fault geometry and half-graben development. Reconstructions of the prefaulting stratigraphy suggest that the footwall is exhumed by ~6 km, but this amount of exhumation is much greater than the estimated thickness of basin fill. Either there was a fluvial outlet for sediment to leave the hanging-wall basin, or much of the exhumation was accomplished by low-angle extensional faulting rather than erosion.

a slip surface in the core zone of the Wasatch fault at Stop 3 (Fig. 8). The slip surface is less than 1 m thick and is composed mostly of polished limestone breccia. Several meters of pulverized carbonate gouge surround the slip surface. Lenses of limestone wall rock are scattered throughout the gouge. The slip surface strikes N-S and dips 42–46°W. It is characterized by a series of steeply dipping corrugations with wavelengths of around a meter, which have slickenside lineations on them. The lineations have a rake of 85–88°S and are up to 160 cm in length, with the majority measuring 30–60 cm in length. Most lineations widen and shallow down dip in a wedge-shaped fashion. Secondary tensile fractures (stretch marks) taper downward into the surface and are mostly horizontal. The longest striae are similar in length to the minimum amount of vertical displacement per event observed in fault trenches across the Provo segment (~1.6–3.3 m; Swan et al., 1980; Machette et al., 1992; Lund and Black, 1998). If the grooves and displacements are from one coseismic event that ruptured the entire Provo segment, it would

Seven Peaks Fault Scarp Excavation for a ski lift, as part of the abandoned Seven Peaks Ski Area at the base of “Y” Mountain, exhumed part of

Figure 7. Fault drag along the Provo segment (between white arrows). Resistant layers of Mississippian limestone are folded down into the fault. Immediately adjacent to the fault, intervening shale between the limestone layers has been removed from between the resistant units and incorporated into the gouge zone. Mount Timpanogos rises in the distance above a strike valley formed by erosion of the Manning Canyon Shale.

Figure 8. Slip surface and architecture of the Seven Peaks fault scarp at Stop 3. Slip surface is polished carbonate-clast breccia with dip-parallel slickenlines. Material in the distance consists of fault gouge (dark gray), blocks of limestone (light gray), and displaced Lake Bonneville deposits (bedded gray units at base).

Neotectonics and paleoseismology of the Wasatch fault have produced an earthquake with a MW ~7.1 (Chang and Smith, 2002). Lee and Bruhn (1996) measured the roughness of the fault surface at this site and concluded that the topography of the fault surface was partly controlled by processes associated with coseismic slip of 1–2 m per earthquake. Stop 4: Salt Lake–Provo Segment Boundary at Corner Creek Travel Directions to Stop 4 Return to I-15 from Stop 3 and proceed northward 22 mi to Exit 288 for Utah Hwy 140. Continue for 0.5 mi to the intersection with Highland Drive. Turn right heading north and east on Highland for 2.7 mi to the intersection of Rambling Road. Turn right onto Rambling Road and proceed up the hill for 0.9 mi, passing through the roundabout onto Sage Hollow Drive. Continue east and uphill on Sage Hollow and onto Gray Fox Drive for 0.5 mi past the roundabout. Turn left and downhill onto Coyote Hollow Drive and park in the designated area. Walk east along the creek to exposures of the Wasatch fault. Geologic Description The boundary between the Salt Lake and Provo segments is a 7-km-wide jog formed by an east-trending cross fault (Fig. 9; Gilbert, 1928; Schwartz and Coppersmith, 1984; Bruhn et al., 1992). At this locality we will observe remnants of the Wasatch fault that formed at a depth of at least 11 km and at temperatures of 350–400 °C (Parry and Bruhn, 1987). The rocks are a mixture of phyllonite (mica-rich mylonite), cataclasite, and thin bands of relict pseudotachylite (frictional melt rock) that provide evidence

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for processes of fault creep and earthquake rupturing in the hypocentral regions of magnitude ~7 earthquakes (Parry and Bruhn, 1986; Bruhn et al., 1990; Yonkee and Bruhn, 1990). The exhumed fault rock forms a thin carapace on the margins of the Little Cottonwood stock, an Oligocene granitic intrusion that is mostly quartz monzonite (Fig. 9). The fault segment boundary defines a roughly cylindrical structure plunging 20°– 25° W-SW beneath the eastern Traverse Mountains, a bedrock ridge of Paleozoic sedimentary rocks and Tertiary volcanic rocks that extends outward from the mountain front in the hanging wall of the Wasatch fault (Fig. 9). Fault slip directions are mostly west to west-southwest based on the orientation of slickenlines and angular relations between foliation and shear surfaces preserved in the phyllonite. Although minor faults are highly variable in orientation and exhibit a broad range of slip directions, the average slip directions for the larger fault sections are consistent with a best-fit regional stress tensor having a maximum principal stress (σ1) plunging steeply SW, a gently NNW-SSE–plunging σ2 axis, and gently ENE-WSW–plunging σ3 axis (Yonkee and Bruhn, 1990). Cross sections constructed through the fault carapace reveal several distinctive zones of deformation and alteration (Fig. 10). Granite in the footwall (zone 0) is cut by widely spaced fractures that form sets of steeply dipping and subhorizontal joints, and has undergone little hydrothermal alteration. This rock grades upward into a transition zone of heterogeneously deformed and altered rock with thickness that varies from ~20 m to >200 m. The lower, or zone 1, part of this transition zone is dissected by decimeter-spaced fractures and an increase in the number

Figure 9. Map of the Salt Lake–Provo segment boundary in the Wasatch fault zone, showing (A) fault traces and location of Little Cottonwood Stock, and (B) segment boundary geology and spatial distribution of fault-zone and country rock. Map symbols 0, 1, 2 and 3 refer to rock zones described in the text. Cz–Pz—Cenozoic and Paleozoic rocks; Q—undifferentiated Quaternary deposits. The Wasatch fault trace is mapped as a thick and double hatched line. Reproduced from Evans et al. (1997) and based on work by Yonkee and Bruhn (1990).

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Figure 10. Structural cross sections of the Wasatch fault along lines A–A′ and B–B′ shown in Figure 9. The numbers on the variably shaded units refer to the fault rock zones discussed in the text. Block diagrams to the right side of the figure are sketches illustrating the fracture and vein patters observed in the various fault rock zones. Modified from Evans et al. (1997) and based on work by Yonkee and Bruhn (1990).

of minor faults compared to zone 0. The upper, more deformed part of the transition zone (zone 2) is characterized by variably oriented fractures with centimeter-scale spacing and many small faults. Quartz veins are common throughout this zone, most of the veins dip steeply, and many veins display evidence for repeated cracking and resealing. The quartz veins reflect episodic dilation and horizontal extension within the fault zone. The transition zone changes upward into the fault slip zone (zone 3) that is generally <10 m thick and is where most displacement occurred based on the presence of large slip surfaces and the intensity of deformation. Rocks in this zone include fault breccia, cataclasite, phyllonite, and large striated fault surfaces. Veins of partly altered pseudotachylite both cross cut and are deformed by the phyllonitic (ductile) and cataclastic (brittle) fabrics, recording alternating periods of quasiplastic creep and rapid brittle faulting, presumably caused by rupture propagation and frictional sliding and heating during earthquakes (Yonkee and Bruhn, 1990). Cataclasite dikes that were injected outward from fault surfaces also provide evidence of very rapid fracturing and material transport. The intimate relation between hydrothermal and deformation processes within the fault zone is perhaps the most important observation to make at this locality. Hydrothermal alteration is concentrated along fracture and vein networks in the transition zone and pervades the fault-slip zone. There are two alteration mineral assemblages: (1) the oldest and most widespread is a quartz-muscovite-chlorite-epidote assemblage that formed at T >300–350 °C, and (2) a younger and spatially restricted assemblage of prehnite and clay minerals that formed at T <200 °C within the upper few km of the fault zone. Fluid inclusion data indicate transient changes in paleofluid pressure between lithostatic and hydrostatic that was presumably associated with earthquake cycles while the rocks were at depths of ~10 km (Parry

and Bruhn, 1986). The evidence suggests that fracturing and hydrothermal alteration of feldspar to micaceous assemblages at temperatures of 350–400 °C dramatically weakened the granitic protolith, enhanced ductile deformation, and consequently localized shearing within the fault zone while the unaltered granite in the footwall remained undeformed. Granitic rocks located only tens of meters from the hydrothermal system were largely unaffected by chemically reactive fluids in the fault zone and by earthquake ruptures that propagated only tens of meters away. Stop 5: The Wasatch Fault at Little Cottonwood Canyon Travel Directions to Stop 5 Retrace route from Stop 4 to Highland Drive. Turn NE on Highland and proceed to the intersection of 1300 East. Turn north on 1300 East and continue to the Draper Parkway. Turn right (E) on Draper Pkwy, which will become 1700 East. Continue north on 1700 East to the intersection with Wasatch Boulevard. Turn east onto Wasatch Boulevard and continue east and then north to the park at the intersection of Wasatch Blvd and Little Cottonwood Canyon Road. Turn into the park for Stop 5. To return to the field trip start (Salt Palace Convention Center), proceed west on Little Cottonwood Canyon Road, which will become 9000 South, and enter I-15 driving north toward Salt Lake City. Proceed north for 11 mi to Exit 306 (600 South). After exiting, proceed east on 600 South and then turn left (N) on West Temple. Continue north for ~0.75 mi to the convention center. Introduction At the mouth of Little Cottonwood Canyon, a prominent glacier-carved valley in the Wasatch Range, the Wasatch fault is expressed by some of the most spectacular and complex scarps

Neotectonics and paleoseismology of the Wasatch fault found anywhere along its length (Fig. 11). The fault zone is up to 400 m wide and consists of prominent curvilinear and en echelon west-facing and antithetic scarps that form a major graben in glacial moraines, Lake Bonneville deposits, and alluvium (Fig. 12). Individual west-facing scarps within the zone reach heights of 35–40 m and antithetic scarps range in height from 7 to 20 m (Schwartz and Lund, 1988). Little Cottonwood Canyon is on the Salt Lake City segment of the Wasatch fault (Fig. 1), one of ten seismogenically independent fault segments identified by Machette et al. (1992) that are distinguished by their unique paleoearthquake chronologies. The Salt Lake City segment is the most heavily urbanized Wasatch fault segment. Urbanization began at the north end of the segment in 1847 with the founding of Salt Lake City and has moved progressively southward over the ensuing 158 years. As a result, most fault exposures on the northern part of the segment were either destroyed or made inaccessible long before paleoseismology developed as a science. In the late 1970s when geologists first began conducting modern paleoseismic investigations on the Salt Lake City segment, they were forced by circumstances to look for sites on the southern part of the segment. However, the fault scarps at Little Cottonwood Canyon are so spectacular and instructive that they drew the attention of G.K. Gilbert, one of America’s preeminent early geologists, nearly 130 years ago, thus securing for the Wasatch fault at Little Cottonwood Canyon a prominent place in the geologic history of the Basin and Range. G.K. Gilbert G.K. Gilbert (1843–1918) was one of the most perceptive geologists ever to work in the American West. His regional investigations of Basin and Range geology with the Wheeler Survey (1871–1874), Powell Survey (1875–1879), and U.S. Geological Survey (1879–1883) resulted in a number of scientific firsts and benchmark studies that are classics of American geologic thought. Many of his theories have withstood the test of time, notably his contributions to the understanding of mountain building processes and earthquakes in the Basin and Range Province. He was the first to recognize that faulting and not folding is the primary mechanism responsible for mountain building in the interior basins of Utah and Nevada (Gilbert, 1872, 1875). He was also the first to recognize “piedmont” scarps as evidence that the mountains are the result of incremental movement along rangebounding faults during earthquakes (Gilbert, 1875, 1890, 1928). The fault scarps at Little Cottonwood Canyon occupy a prominent place in the development of G.K. Gilbert’s theories concerning mountain building and earthquakes, and his field notes contain numerous sketches of the scarps and moraines at this location (Wallace, 1980; Hunt, 1982). Gilbert visited this locality twice in 1877 to make notes on the geology and surface-water resources and again in 1880 to spend several days mapping the geology. Gilbert (1890) mapped and profiled the glacial moraines, and mapped the fault scarps where they cross the moraines and fan deposits defining a major graben. He noted the complexity of the fault zone and that it contains opposing

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scarps. Based on his observations that the scarps in the moraines are higher than scarps in younger alluvial deposits, he concluded that “it is evident that the total displacement was accomplished by a series of efforts” (Gilbert, 1890, p. 347). In 1883, confident that his theories about mountain building and earthquakes were correct, Gilbert issued an earthquake hazard warning to the residents of Salt Lake City. In an article in the Salt Lake City Tribune (20 September 1883) and later reprinted in the American Journal of Science (Gilbert, 1884), he summarized his ideas and emphasized their practical application to Utah. Gilbert’s earthquake warning was remarkable for its recognition that earthquakes in the Basin and Range are unevenly distributed in space and time. He concluded that once an earthquake occurs at a particular location on a fault, it is unlikely that another will take place there until enough time has elapsed for sufficient strain to reaccumulate. He interpreted the presence of young scarps at this location, but their absence farther north, as evidence of long quiescence and strain accumulation in the Salt Lake City area, thus making that subdued section of the fault a prime candidate for a future earthquake. It was not until almost 100 years later, in 1979, that geologists pursuing the budding science of paleoseismology finally returned to Little Cottonwood Canyon to begin the arduous process of quantifying the past behavior and earthquake potential of the Wasatch fault. Modern Paleoseismology Studies In addition to figuring prominently in development of our early understanding of Basin and Range faulting, the Wasatch fault was also the first normal-slip fault in the world to be investigated using modern paleoseismic techniques. Much of what is now considered state-of-the-art paleoseismic practice for investigating faults was originally developed and refined at Little Cottonwood Canyon and at other sites along the Wasatch fault. Woodward-Clyde consultants—the first paleoseismic study. Woodward-Clyde Consultants conducted the first paleoseismic study on the Salt Lake City segment in 1979. Their study was one of the first ever conducted on a normal-slip fault anywhere, and was in the Little Cottonwood graben ~200 m northeast of where we are now standing (Fig. 12). The study included excavation of four trenches and topographic profiling of the faulted Bells Canyon lateral moraine (immediately south and east of our present location). Based on stratigraphic relations in the trenches and the results of early accelerator–mass spectrometry 14C dates, Woodward-Clyde concluded that there was evidence for at least two surface faulting earthquakes during the Holocene at Little Cottonwood Canyon: a penultimate earthquake that occurred shortly before 8–9 ka, and the most recent earthquake that is younger than 8–9 ka, but the timing of which could not be further constrained (Swan et al., 1981; Schwartz and Coppersmith, 1984). A critical result of the Woodward-Clyde study was the recognition of a distinct Salt Lake City segment of the Wasatch fault (Schwartz and Coppersmith, 1984). Utah Geological Survey—South Fork Dry Creek and Dry Gulch Trench sites. Recognizing that it was unlikely that the

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Figure 11. Low sun–angle aerial photograph showing the complexity and width of the Wasatch fault (white arrows) at the mouths of Little Cottonwood and Bells canyons and the location of Stop 5.

Neotectonics and paleoseismology of the Wasatch fault

Figure 12. Surficial geologic map of the mouth of Little Cottonwood Canyon showing the Little Cottonwood megatrench site (LC) and Woodward-Clyde trenches (LC-1–4). Map units: al1, al2, and alp—stream alluvium; af2 and af4—alluvial fan deposits; chs and ca—colluvial and alluvial deposits; gbct and gbco—glacial deposits; lbg—Lake Bonneville lacustrine deposits. Modified from Personius and Scott (1992).

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Woodward-Clyde paleoseismic investigation had determined a complete Holocene earthquake chronology for the Salt Lake City segment, the Utah Geological Survey, with assistance from the U.S. Geological Survey and funding from the National Earthquake Hazards Reduction Program (NEHRP), undertook a series

TABLE 3. TIMING OF SURFACE-FAULTING EARTHQUAKES ON THE SALT LAKE CITY SEGMENT DETERMINED AT THE SOUTH FORK DRY CREEK AND DRY GULCH TRENCH SITES Event

Minimum timing (cal yr B.P.)

Preferred timing (cal yr B.P.)

Maximum timing (cal yr B.P.)

Z

1100

1300

1550

Y

2100

2450

2800

X

3500

3950

4500

W

4950

5300

5750

Note: From Black et al. (1996). Earthquake timing is based on a combination of apparent mean resident time and accelerator mass spectrometer radiocarbon ages obtained from the South Fork Dry Creek and Dry Gulch trenches.

Figure 13. View to the east of the “megatrench” at Little Cottonwood Canyon during the summer of 2000 that exposed 18 m of Holocene and late Pleistocene stratigraphy across two fault scarps. White arrows indicate location of major west-dipping (upper arrow) and east-dipping (lower arrow) fault zones exposed in the trench (McCalpin, 2002). Photo by B. Black.

of paleoseismic studies in the 1980s and 1990s at South Fork Dry Creek and Dry Gulch (Black et al., 1996) ~3 km south of Stop 5. The advantages of the South Fork Dry Creek and Dry Gulch sites were twofold: first, the sites were in nearly pristine condition; second, slip across the Wasatch fault was distributed across a zone roughly 400 m wide on several subparallel scarps whose heights were measured in meters rather than tens of meters, making the scarps easier to trench. The Utah Geological Survey excavated nine trenches across the fault scarps at South Fork Dry Creek and logged a consultant’s trench excavated across one fault scarp at Dry Gulch. Trenching at South Fork Dry Creek and Dry Gulch showed that four rather than two surface-faulting earthquakes have occurred on the Salt Lake City segment in the past ~6000 yr (Table 3), thus greatly reducing the mean recurrence time for surface faulting for the segment during the Holocene and substantially raising the segment’s hazard potential. GEO-HAZ Consulting Inc. and the return to Little Cottonwood Canyon. In 2000, GEO-HAZ Consulting, Inc., also with NEHRP funding, reoccupied the Woodward-Clyde site in the Little Cottonwood graben and excavated a single “megatrench” across two fault scarps totaling 18 m high (McCalpin, 2002). The trench and an accompanying auger hole exposed 26 m of vertical section, roughly four times that of a typical trench on the Wasatch fault zone (Fig. 13). Each of the two fault scarps was underlain by a major, down-to-the-west normal fault with 7–9.5 m of vertical displacement measured on Lake Bonneville lacustrine sediments. The trench also exposed two minor antithetic faults; however, they lacked any surface expression, having been buried by wash-facies colluvium shed from the larger scarps. The megatrench was 65 m long but did not extend across the graben and therefore did not expose the large graben-bounding antithetic fault on the west side of Wasatch Boulevard (Fig. 12). The megatrench contained evidence for seven paleoearthquakes younger than the Bonneville flood at ca. 18 ka and possibly an eighth earthquake that occurred while Lake Bonneville was at or near its highstand at the Bonneville shoreline (Fig. 14). The GEO-HAZ study established the composite surface faulting chronology for the Wasatch fault at Little Cottonwood Canyon shown in Table 4. McCalpin (2002) states that the age constraints for paleoearthquakes resulting from the megatrench study are not as closely limiting as those resulting from the Black et al. (1996) study at the South Fork Dry Creek and Dry Gulch sites (Table 3). He attributes the discrepancy in part to the fact that organic matter in the megatrench was not always found near paleoearthquake event horizons and that some earthquake ages from the megatrench are based on 14C ages on organics from crack fills rather than from paleosols beneath colluvial wedges. Utah Quaternary Fault Parameters Working Group. In 2003, the Utah Geological Survey convened the Utah Quaternary Fault Parameters Working Group, a panel of expert paleoseismologists and seismologists, to make a comprehensive evaluation of the paleoseismic-trenching data available for Utah’s Quaternary

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TABLE 4. EARTHQUAKE TIMING DETERMINED FROM THE MEGATRENCH EXCAVATED AT THE LITTLE COTTONWOOD CANYON SITE Event*

Timing

Z

1.3 ka

Y

2.3 ka

X

3.5 ka

W

5.3 ka

V

7.5 ka

U

9 ka

T

17 ka

S(?)

Figure 14. Hydrograph showing the timing of major shorelines of Pleistocene Lake Bonneville (after Oviatt, 1997).

faults, and where the data permitted, to assign consensus earthquake timing, recurrence interval, and vertical slip-rate estimates for the faults and/or fault sections under review (Lund, 2005). As part of the review, the Working Group considered all of the paleoseismic trenching data available for the Salt Lake City segment. The Working Group concluded that all reliable paleoearthquake timing, recurrence, size, vertical net displacement, and slip-rate information for the Salt Lake City segment comes from studies either conducted at Little Cottonwood Canyon or at the South Fork Dry Creek and Dry Gulch sites a few km to the south (see above). The Working Group’s consensus paleoseismic parameters for the Salt Lake City segment are summarized in Table 5. The Working Group’s consensus earthquake timing and recurrence interval and vertical slip-rate estimates represent the best available paleoseismic information for the Salt Lake City segment until superseded by data from future studies. Summary The Wasatch fault at Little Cottonwood Canyon has been of great interest to geologists for almost 130 yr. This location played a key role both in the initial understanding of Basin and Range tectonics and later in developing the science of paleoseismology. Investigators conducting multiple paleoseismic studies at and near Little Cottonwood Canyon over the past 26 years have identified four surface faulting earthquakes since the middle Holocene (ca. 6 ka) based on colluvial-wedge stratigraphy, tectonic crack fills, and fault terminations. The timing of these earthquakes is provided by 14C ages on both charcoal and bulk organic samples (primarily paleosol A horizons and tectonic crack-fill deposits). Both the number of earthquakes and their timing are considered well constrained, as is the resulting mean recurrence

Possibly between 17 and 20 ka

Note: Earthquake timing based on McCalpin (2002). *Events Z to W: identification and timing based on colluvial-wedge stratigraphy and radiocarbon ages. Events V to S: identification based on stratigraphic relations in the trench that include the absence or thinning of deposits attributed to erosion following surface faulting and retrodeformation of faulting. The timing of events V and U are constrained by sparse 14C ages; the timing of events T and S are inferred from Lake Bonneville stratigraphy at the trench site.

between surface-faulting earthquakes since the middle Holocene of 1300 ± 400 yr (Lund, 2005). Three older earthquakes (T, U, V), and a possible fourth oldest earthquake (S) were identified from secondary stratigraphic relations in the Little Cottonwood megatrench and retrodeformation of faulting in the trench. These four earthquakes lack direct stratigraphic or structural evidence (colluvial wedges, crack-fill deposits, fault terminations) of their occurrence (McCalpin, 2002). However, even though the timing of these earthquakes is only broadly constrained, a comprehensive review by the Utah Quaternary Fault Parameters Working Group found the evidence for these earthquakes compelling (Lund, 2005). Available data indicate two recurrence intervals of ~2 k.y. in mid- to early Holocene time, preceded by a >7 k.y. period of surface faulting quiescence. Evidence from Little Cottonwood Canyon and other nearby sites shows that the elapsed time since the most recent surface faulting on the Salt Lake City segment is equal to or greater than the Working Group’s preferred recurrence interval estimate (500–1300–2400 yr) for the segment, indicating that the Salt Lake City segment is a candidate for the next large surface faulting earthquake on the Wasatch fault. CONCLUSIONS The Wasatch fault poses a significant seismic hazard to north-central Utah’s urban population, based on decades of research that has identified a well-documented chronology of Holocene surface faulting along the fault’s six central segments. Study of the Wasatch fault has contributed greatly to the development of the science of paleoseismology and to our

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TABLE 5. UTAH QUATERNARY FAULT PARAMETERS WORKING GROUP CONSENSUS PALEOSEISMIC PARAMETERS FOR THE SALT LAKE CITY SEGMENT OF THE WASATCH FAULT Consensus earthquake timing*

Interevent recurrence†

Weighted mean recurrence three most recent events§

Consensus holocene mean recurrence#

Consensus vertical slip rate**

500–1300–2400 yr

0.6–1.2–4.0 mm/yr

Z: 1300 ± 650 cal yr B.P. 1150 ± 900 cal yr Y: 2450 ± 550 cal yr B.P. 1500 ± 800 cal yr

1300 ± 400 cal yr

X: 3950 ± 550 cal yr B.P. 1350 ± 900 cal yr W: 5300 ± 750 cal yr B.P. V: ca. 7.5 ka (<8.8–9.1 ka; >5.1–5.3 ka) U: ca. 9 ka (≤ 9.5-9.9 ka) T: ca. 17 ka S(?): 17–20 ka

Note: From Lund (2005). *Timing for earthquakes W, X, Y, and Z is from Black et al. (1996). The ± confidence limits have been increased for each earthquake to accommodate the full range of limiting 14C ages used to constrain the timing of the earthquakes; the resulting ± ranges account for both the laboratory and geologic uncertainty associated with the timing of each earthquake (Lund, 2005). McCalpin (2002) identified earthquakes V, U, T, and S at Little Cottonwood Canyon based on stratigraphic relations and a retrodeformation analysis of structural relations in the megatrench. No direct evidence (colluvial wedges, tectonic crack fills, fault terminations) was found to document these earthquakes, and consequently their timing is only broadly constrained. † The ± confidence limits equal the square root of the sum of the squares of the individual ± confidence limits for each bracketing earthquake rounded to nearest 100 yr. § Weighted mean rounded to nearest 100 yr; 2-sigma confidence limits rounded to nearest 100 yr. The weighted mean recurrence for the three most recent surface-faulting earthquakes on the Salt Lake City Segment (SLCS) is ≤ to the elapsed time since the most recent surface-faulting earthquake on the SLCS. # The Working Group’s preferred recurrence-interval and confidence-limit estimates (bold) for the SLCS are based on currently available information on earthquake timing and variability for the SLCS (Lund, 2005). **The Working Group’s preferred vertical slip-rate and confidence-limit estimates (bold) for the SLCS are based on currently available information on earthquake timing and displacement (Lund, 2005).

understanding of Basin and Range Province fault mechanics and tectonic geomorphology. In traversing the Wasatch fault from Nephi to Salt Lake City, we have highlighted evidence for earthquake generation in the subsurface based on the structure and rheology of fault-zone rocks, and identified the timing, displacement, and extent of Holocene and late Pleistocene surface faulting along the fault’s central segments. Collectively, the field-trip stops have allowed us to peer into the Wasatch fault’s “earthquake engine” and to consider the natural complexity of its pulse. Along the fault, vertically displaced alluvial, glacial, and lacustrine deposits demonstrate a long history of Quaternary normal faulting; however, geomorphic relations and paleoseismic trench studies on the central Wasatch fault segments suggest higher rates of faulting during the Holocene than during the late Pleistocene. Additional paleoseismic studies along the Wasatch fault are necessary to test possible slip-rate variations and segmentation along the fault, and to refine the timing, recurrence,

and coseismic displacement of surface faulting earthquakes. Important unresolved issues remain, including determining the fault’s subsurface geometry and potential for aseismic creep and resolving discrepancies between geologic slip rates and geodetic extension rates. Evaluating these issues is critical to characterize future surface faulting earthquakes on the Wasatch fault and the earthquake hazard along the Wasatch Front. ACKNOWLEDGMENTS We thank Joel Pederson (Utah State University), and Gary Christenson and Michael Hylland (Utah Geological Survey) for carefully reviewing this field trip guide. Research on the Nephi segment (Stop 1) was supported by National Science Foundation grant EAR-0207373; fault trenching at the Santaquin site (Stop 2) was supported by the Utah Geological Survey and the U.S. Geological Survey National Earthquake Hazards Reduction Program, and land access was granted by the U.S. Forest Service.

Neotectonics and paleoseismology of the Wasatch fault REFERENCES CITED Andrews, D.J., and Bucknam, R.C., 1987, Fitting degradation of shoreline scarps by a nonlinear diffusion model: Journal of Geophysical Research, v. 92, no. B12, p. 12,857–12,867. Arabasz, W.J., and Julander, D.R., 1986, Geometry of seismically active faults and crustal deformation within the Basin and Range—Colorado Plateau transition, in Mayer, L., ed., Extensional tectonics of the southwestern United States: A perspective on processes and kinematics: Geological Society of America Special Paper 208, p. 43–74. Arabasz, W.J., Smith, R.B., and Richins, W.D., 1979, Earthquake studies along the Wasatch Front, Utah: Network monitoring, seismicity and seismic hazards, in Arabasz, W.J., Smith, R.B., and Richins, W.D., eds., Earthquake studies in Utah, 1850–1978: Salt Lake City, University of Utah Seismograph Stations, p. 253–285. Arabasz, W.J., Pechmann, J.C., and Brown, E.D., 1992, Observational seismology and the evaluations of earthquake hazards and risk in the Wasatch Front area, in Gori, P.L., and Hays, W.W., eds., Assessment of regional earthquake hazards and risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1500-D, 36 p. Armstrong, P.A., Ehlers, T.A., Chapman, D.S., Farley, K.A., and Kamp, P.J.J., 2003, Exhumation of the central Wasatch Mountains, 1: Pattern and timing deduced from low-temperature thermochronology data: Journal of Geophysical Research, v. 108, no. B3, p. 2172, doi: 10.1029/2001JB001708. Armstrong, P.A., Taylor, A.R., and Ehlers, T.A., 2004, Is the Wasatch fault (Utah, United States) footwall segmented over million-year time scales?: Geology, v. 32, p. 385–388, doi: 10.1130/G20421.1. Black, B.D., Lund, W.R., Schwartz, D.P., Gill, H.E., and Mayes, B.H., 1996, Paleoseismic investigation on the Salt Lake City segment of the Wasatch fault zone at the South Fork Dry Creek and Dry Gulch sites, Salt Lake County, Utah: Utah Geological Survey Special Study 92, 22 p. Black, B.D., Hecker, S., Hylland, M.D., Christenson, G.E., and McDonald, G.N., 2003, Quaternary fault and fold database and map of Utah: Utah Geological Survey Map 193DM, scale 1:500,000, CD-ROM. Bruhn, R.L., Yonkee, W.A., and Parry, W.T., 1990, Structural and fluid-chemical properties of seismogenic normal faults: Tectonophysics, v. 175, p. 139–157, doi: 10.1016/0040-1951(90)90135-U. Bruhn, R.L., Gibler, P.R., Houghton, W., and Parry, W.T., 1992, Structure of the Salt Lake segment, Wasatch normal fault zone: Implications for rupture propagation during normal faulting, in Gori, P.L., and Hays, W.W., eds., Assessment of regional earthquake hazards and risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1500-C, 16 p. Bruhn, R.L., Parry, W.T., Yonkee, W.A., and Thompson, T., 1994, Fracturing and hydrothermal alteration in normal fault zones: Pure and Applied Geophysics, v. 142, p. 609–643, doi: 10.1007/BF00876057. Bryant, B., Naeser, C.W., Marvin, R.F., and Mehnert, H.H., 1989, Ages of late Paleogene and Neogene tuffs and the beginning of rapid regional extension, eastern boundary of the Basin and Range Province near Salt Lake City, Utah: U.S. Geological Survey Bulletin 1787, 37 p. Bucknam, R.C., 1978, Northwestern Utah seismotectonics studies, in Seiders, W., and Thomson, J., compilers, Summaries of technical reports, v. VII: Menlo Park, California, U.S. Geological Survey Office of Earthquake Studies, p. 64. Chang, W.L., 1998, Earthquake hazards on the Wasatch fault—tectonically induced flooding and stress triggering of earthquakes [M.S. thesis]: Salt Lake City, University of Utah, 123 p. Chang, W.L., and Smith, R.B., 2002, Integrated seismic-hazard analysis of the Wasatch Front, Utah: Bulletin of the Seismological Society of America, v. 92, no. 5, p. 1904–1922, doi: 10.1785/0120010181. DuRoss, C.B., 2004, Spatial and temporal trends of surface rupturing on the Nephi segment of the Wasatch fault, Utah—implications for fault segmentation and the recurrence of paleoearthquakes [M.S. thesis]: Salt Lake City, University of Utah, 120 p. DuRoss, C.B., and Bruhn, R.L., 2005, Active tectonics of the Nephi segment, Wasatch fault zone, Utah, in Lund, W.R., ed., Western States Seismic Policy Council Proceedings Volume of the Basin and Range Province Seismic Hazards Summit II: Utah Geological Survey Miscellaneous Publication 05-2, CD-ROM. Emmons, B.F., 1878, U.S. Geological Expedition 40th Parallel Report, v. 2, 357 p. Evans, J.P., Yonkee, W.A., Parry, W.T., and Bruhn, R.L., 1997, Fault-related rocks of the Wasatch normal fault: Brigham Young University Geology Studies, v. 42, Part II, p. 279–297.

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Evans, S.H., Parry, W.T., and Bruhn, R.L., 1985, Thermal, mechanical and chemical history of Wasatch fault cataclasite and phyllonite, Traverse Mountains area, Salt Lake City, Utah: From K/Ar and fission track measurements: U.S. Geological Survey Open-File Report 86-31, p. 410–415. Friedrich, A.M., Wernicke, B.P., Niemi, N.A., Bennett, R.A., and Davis, J.L., 2003, Comparison of geodetic and geologic data from the Wasatch region, Utah, and implications for the spectral character of Earth deformation at periods of 10 to 10 million years: Journal of Geophysical Research, v. 108, no. B4, 2199, doi: 10.1029/2001JB000682. Gilbert, G.K., 1872, Reports on exploration in Nevada and Arizona [abs.]: Forty-second Congress, second session, Senate Document 65, p. 90–94. Gilbert, G.K., 1875, Report on the geology of portions of Nevada, Utah, California, and Arizona examined in the years 1871 and 1872: Report on U.S. Geographical and Geological Surveys west of the 100th meridian, v. 3, Geology: Pt. 1, p. 17–187. Gilbert, G.K., 1884, A theory of the earthquakes of the Great Basin, with a practical application: American Journal of Science, Third Series, v. XXXVII, no. 157, Article XI, p. 49–53. Gilbert, G.K., 1890, Lake Bonneville: U.S. Geological Survey Monograph 1, 438 p. Gilbert, G.K., 1928, Studies of Basin and Range structure: U.S. Geological Survey Professional Paper 153, 89 p. Hanson, K.L., Swan, F.H., and Schwartz, D.P., 1981, Study of earthquake recurrence intervals on the Wasatch fault, Utah: San Francisco, California, Woodward-Clyde Consultants, Sixth-annual technical report prepared for U.S. Geological Survey under contract no. 14-08-0001-19115, 22 p. Harris, R.A., Smith, R.B., Chang, W.L., Meertens, C., and Friedrich, A., 2000, Temporal distribution of extensional strain across the southern Wasatch fault zone—geological constraints for the GPS velocity field: Eos (Transactions, American Geophysical Union), v. 81, p. 1230. Harty, K.M., Mulvey, W.E., and Machette, M.N., 1997, Surficial geologic map of the Nephi segment of the Wasatch fault zone, eastern Juab County, Utah: Utah Geological Survey Map 170, 14 p., 1 pl., scale 1:50,000. Hintze, 1988, Geologic history of Utah: Provo, Brigham Young University, 202 p. Hunt, C.B., ed., 1982, Pleistocene Lake Bonneville, ancestral Great Salt Lake, as described in notebooks of G.K. Gilbert, 1875–1880: Brigham Young Geology Studies, v. 29, part 1, 231 p. Hylland, M.D., and Machette, M.N., 2004, Interim surficial geologic map of the Levan segment of the Wasatch fault zone, Juab and Sanpete Counties, Utah, in Christenson, G.E., Ashland, F.X., Hylland, M.D., McDonald, G.N., and Case, B., Database compilation, coordination of earthquakehazards mapping, and study of the Wasatch fault and earthquake-induced landslides, Wasatch Front, Utah: U.S. Geological Survey, National Earthquake Hazards Reduction Program Final Technical Report, award no. 03HQAG0008, variously paginated, scale 1:50,000. Jackson, M., 1991, Number and timing of Holocene paleoseismic events on the Nephi and Levan segments, Wasatch fault zone, Utah: Utah Geological Survey Special Study 78, 23 p. King, C., 1878, U.S. Geological Expedition 40th Parallel Report, v. 1, p. 745–746. Kowallis, B.J., Ferguson, J., and Jorgensen, G.J., 1990, Uplift along the Salt Lake segment of the Wasatch fault from apatite and zircon dating in the Little Cottonwood stock, in Durani, S.A., and Benton, E.V., eds., Proceedings of the 6th International Fission Track Dating Workshop, Nuclear Tracks and Radiation Measurements 17: Oxford, Pergamon, p. 325–329. Lee, J.-J., and Bruhn, R.L., 1996, Characterization of the structural anisotropy of normal fault surfaces: Journal of Structural Geology, v. 18, p. 1043– 1059, doi: 10.1016/0191-8141(96)00022-3. Lund, W.R., 2005, Consensus preferred recurrence-interval and vertical sliprate estimates—Review of Utah paleoseismic-trenching data by the Utah Quaternary Fault Parameters Working Group: Utah Geological Survey Bulletin 130, CD-ROM. Lund, W.R., and Black, B.D., 1998, Paleoseismic investigation at Rock Canyon, Provo segment, Wasatch fault zone, Utah County, Utah: Special Study #93, Utah Geological Survey, 21 p. Lund, W.R., Schwartz, D.P., Mulvey, W.E., Budding, K.E., and Black, B.D., 1991, Fault behavior and earthquake recurrence of the Provo segment of the Wasatch fault zone at Mapleton, Utah County, Utah: Utah Geological and Mineral Survey Special Studies 75, 41 p. Mabey, D.R., 1992, Subsurface geology along the Wasatch fault area, Utah, in Gori, P.L., and Hays, W.W., eds., Assessment of regional earthquake hazards and risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1500-C, 16 p.

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Machette, M.N., 1992, Surficial geologic map of the Wasatch fault zone, eastern part of Utah Valley, Utah County and parts of Salt Lake and Juab Counties, Utah: U.S. Geological Survey Miscellaneous Investigations Series Map I-2095, scale 1:50,000, pamphlet 30 p. Machette, M.N., Personius, S.F., and Nelson, A.R., 1992, Paleoseismology of the Wasatch fault zone—A summary of recent investigations, interpretations, and conclusions, in Gori, P.L., and Hays, W.W., eds., Assessment of regional earthquake hazards and risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1500-A, p. A1–A71. Martinez, L.J., Meertens, C.M., and Smith, R.B., 1998, Rapid deformation rates along the Wasatch fault zone, Utah, from first GPS measurements with implications for earthquake hazard: Geophysical Research Letters, v. 25, no. 4, p. 567–570, doi: 10.1029/98GL00090. Mattson, A., and Bruhn, R.L., 2001, Fault slip rates and initiation age based on diffusion equation modeling—Wasatch fault zone and eastern Great Basin: Journal of Geophysical Research, v. 106, no. B7, p. 13,739– 13,750, doi: 10.1029/2001JB900003. McCalpin, J.P., editor, 1996, Paleoseismology: San Diego, Academic Press, 588 p. McCalpin, J.P., 2002, Post-Bonneville paleoearthquake chronology of the Salt Lake City segment, Wasatch fault zone, from the 1999 “Megatrench” site: Utah Geological Survey Miscellaneous Publication 02-7, 37 p. McCalpin, J.P., and Nishenko, S.P., 1996, Holocene paleoseismicity, temporal clustering, and probabilities of future large (M >7) earthquakes on the Wasatch fault zone, Utah: Journal of Geophysical Research, v. 101, no. B3, p. 6233–6253, doi: 10.1029/95JB02851. Oviatt, C.G., 1997, Lake Bonneville fluctuations and global climate change: Geology, v. 25, p. 155–158, doi: 10.1130/0091-7613(1997)025<0155: LBFAGC>2.3.CO;2. Pack, F.J., 1926, New discoveries relating to the Wasatch fault: American Journal of Science, v. 27, p. 399–410. Parry, W.T., and Bruhn, R.L., 1986, Pore fluid and seismogenic characteristics of fault rock at depth on the Wasatch fault, Utah: Journal of Geophysical Research, v. 91, p. 730–744. Parry, W.T., and Bruhn, R.L., 1987, Fluid inclusion evidence for minimum 11 km vertical offset on the Wasatch normal fault, Utah: Geology, v. 15, p. 67–70, doi: 10.1130/0091-7613(1987)15<67:FIEFMK>2.0.CO;2. Parry, W.T., Wilson, P.N., and Bruhn, R.L., 1988, Pore fluid chemistry and chemical reactions on the Wasatch normal fault, Utah: Geochimica et Cosmochimica Acta, v. 52, p. 2053–2063, doi: 10.1016/00167037(88)90184-6. Personius, S.F., and Scott, W.E., 1992, Surficial geology of the Salt Lake City segment and parts of adjacent segments of the Wasatch fault zone, Davis,

Salt Lake, and Utah Counties, Utah: U.S. Geological Survey Miscellaneous Investigation Series Map I-2106, scale 1:50,000. Schwartz, D.P., and Coppersmith, K.J., 1984, Fault behavior and characteristic earthquakes—Examples from the Wasatch and San Andreas fault zones: Journal of Geophysical Research, v. 89, no. B7, p. 5681–5698. Schwartz, D.P., and Lund, W.R., 1988, Paleoseismicity and earthquake recurrence at Little Cottonwood Canyon, Wasatch fault zone, Utah, in Machette, M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range Province, Guidebook for Field Trip Twelve: Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 82–85. Smith, R.B., and Bruhn, R.L., 1984, Intraplate extensional tectonics of the eastern Basin-Range: Inferences on structural style from seismic reflection data, regional geophysics and thermal mechanical models of brittle-ductile deformation: Journal of Geophysical Research, v. 87, p. 5733–5762. Swan, F.H., III, Schwartz, D.P., and Cluff, L.S., 1980, Recurrence of moderate to large magnitude earthquakes produced by surface faulting on the Wasatch fault zone, Utah: Bulletin of the Seismological Society of America, v. 70, p. 1431–1462. Swan, F.H., III, Hanson, K.L., Schwartz, D.P., and Knuepfer, P.L., 1981, Study of earthquake recurrence intervals on the Wasatch fault, Utah—Little Cottonwood Canyon site: U.S. Geological Survey Open-File Report 81-450, 30 p. Wallace, R.E., 1980, G.K. Gilbert’s studies of faults, scarps, and earthquakes, in Yochelson, E.L., ed., The scientific ideas of G.K. Gilbert; An assessment on the occasion of the centennial of the United States Geological Survey (1879–1979): Geological Society of America Special Paper 183, p. 35–44. Wheeler, R.L., and Krystinik, K.B., 1992, Persistent and nonpersistent segmentation of the Wasatch fault zone, Utah–statistical analysis for evaluation of seismic hazard, in Gori, P.L., and Hays, W.W., eds., Assessment of regional earthquake hazards and risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1500-B, p. B-1–B-47. Yonkee, W.A., and Bruhn, R.L., 1990, Geometry and mechanics of a structural boundary, Wasatch fault zone, Utah, in Bruhn, R.L., Lee, J-J, and Yonkee, W.A., eds., Structural properties of the American Fork, Provo, and Spanish Fork subsegments, Wasatch normal fault zone, Utah: Utah Geological and Mineral Survey Open-File Report 186, 50 p. Zoback, M.L., 1983, Structure and Cenozoic tectonism along the Wasatch fault zone, Utah, in Miller, D.M., Todd, V.R., and Howard, K.A., eds., Tectonic and stratigraphic studies in the eastern Great Basin: Geological Society of America Memoir 157, p. 3–27.

Printed in the USA

Geological Society of America Field Guide 6 2005

Pocatello Formation and overlying strata, southeastern Idaho: Snowball Earth diamictites, cap carbonates, and Neoproterozoic isotopic profiles Paul Karl Link Department of Geosciences, Idaho State University, Pocatello, Idaho 83209-8072, USA Frank A. Corsetti Nathaniel J. Lorentz Department of Earth Sciences, University of Southern California, Los Angeles, California 90089-0740, USA

ABSTRACT This one-day field trip examines two Neoproterozoic sections in the Portneuf Narrows area, SE of Pocatello, Idaho. Rocks to be examined on the first traverse belong to the Scout Mountain Member, Pocatello Formation, and include <710 Ma glacial diamictites, dolomite cap, 667 Ma tuff, and upper caplike carbonate. The second traverse, in Blackrock Canyon, exposes the cyclic Blackrock Canyon Limestone, with upward-shallowing siliciclastic to carbonate cycles and microbial mounds. Keywords: Neoproterozoic, Pocatello Formation, Sturtian glaciation, δ13C values, cap carbonate.

INTRODUCTION AND SIGNIFICANCE

upward-shallowing siliciclastic to carbonate cycles and microbial mounds and with neutral to positive δ13C values.

This one-day field trip examines two Neoproterozoic sections in the Portneuf Narrows area, southeast of Pocatello, Idaho. These strata are significant in that they record late Neoproterozoic (Sturtian) glaciation, volcanism, and postglacial marine sedimentation. These rocks have recently yielded carbon isotope and U-Pb geochronologic data that allow tie points to other Neoproterozoic sections. Rocks to be examined on the first traverse belong to the Scout Mountain Member, Pocatello Formation, and include <709 Ma glacial diamictites and dolomite cap with negative δ13C values overlain by an erosional sequence boundary. Above this is an upward fining succession, 667 Ma fluvial reworked tuff, and upper transgressive caplike carbonate, also with negative δ13C values. The second traverse, in Blackrock Canyon, exposes the cyclic Blackrock Canyon Limestone, with

LOGISTICS This trip includes two traverses in the Portneuf Narrows area along the Portneuf River and Interstate Highway 15, just southeast of Pocatello, Idaho. The locations of the stops are shown in Figures 1 and 2. The directions to the stops are as follows. From the Portneuf Area Exit from Interstate 15, south of Pocatello, Idaho, head east on the frontage road to the first road to the north, Blackrock Canyon Road, and go under the freeway. For Stop 1, turn immediately left (west) onto a dirt road through a Bureau of Land Management gate. Drive 0.5 mi (0.8 km) up this road and park in a broad space before the road climbs steeply up the hill to the north. The traverse is high on the hill to the north-

Link, P.K., Corsetti, F.A., and Lorentz, N.J., 2005, Pocatello Formation and overlying strata, southeastern Idaho: Snowball Earth diamictites, cap carbonates, and Neoproterozoic isotopic profiles, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 251–259, doi: 10.1130/2005.fld006(12). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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Figure 1. The Neoproterozoic succession near Pocatello, Idaho, and locality map of the Blackrock Canyon Limestone locality described and sampled for this study. Note that the Gibson Jack Formation is part of the Brigham Group, but it is not included here. Section adapted from Link et al. (1993) and references therein. U-Pb dates are after Fanning and Link (2004). Age of ca. 580 Ma (K-Ar, recalculated) on Browns Hole Formation is from Christie-Blick and Levy (1989). Interregionally important sequence boundaries 1–4 and paleoenvironmental interpretations are after Link et al. (1993).

Pocatello Formation and Overlying Strata, Southeastern Idaho

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Figure 2. Geologic map of the Portneuf Narrows area, from Link (1987). Cross sections along lines shown are in Link (1987). An updated geological map of the Inkom quadrangle is in preparation by D.W. Rodgers and colleagues at Idaho State University.

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west and requires a climb of ~800 ft (245 m). If the Forest Service gate is locked, a walk of an extra 0.5 mi (0.8 km) is required. Stop 2 is on the east side of Blackrock Canyon itself, about one mile (1.7 km) up the road from the highway underpass (Fig. 1). Park the cars on the east side of the road and climb the slope to the east. GEOLOGIC BACKGROUND Neoproterozoic rocks of western North America are typically divided into a lower diamictite and volcanic succession and

an upper terrigenous detrital succession (Stewart and Suczek, 1977). In general, the diamictite-volcanic succession is related to ca. 700 Ma rifting of the North American craton, and the terrigenous-detrital succession represents the transition into the postrift development of a passive margin (e.g., Stewart, 1970, 1982; Stewart and Poole, 1974). In southeastern Idaho (Figs. 1 and 3), the Pocatello Formation represents the diamictite and volcanic succession and the Brigham Group represents the terrigenous-detrital succession. The Blackrock Canyon Limestone is located stratigraphically between the Pocatello Formation and Brigham Group.

Figure 3. Synthesis stratigraphic column for Neoproterozoic rocks of southeastern Idaho with correlation to Death Valley (from Corsetti et al., 2006). Sequence boundaries 1 and 2 are thought to be correlative between the regions, and carbon isotopic profiles are consistent between sequence boundaries 1 and 2. Isotopic tie points a, b, and c are given for clarity. This correlation matches the Death Valley Noonday Dolomite with the first Pocatello Formation cap carbonate and does not favor the proposal that the upper Kingston Peak Formation of Death Valley represents Marinoan glaciation (Prave, 1999). Sections and sequence boundaries are adapted from Link et al. (1993) and references therein. Pocatello area isotopic data from Lorentz et al. (2004) and Smith et al. (1994); Pocatello dates after Fanning and Link (2004); Death Valley isotopic data from Corsetti and Kaufman (2003).

Pocatello Formation and Overlying Strata, Southeastern Idaho Pocatello Formation The Pocatello Formation is divided into three members: the Bannock Volcanic Member, the Scout Mountain Member, and the (informal) upper member (Link, 1983) (Figs. 1, 3, and 4). The base of the Pocatello Formation is not exposed. The Bannock Volcanic Member exists as a gradational, lenticular body within the lower Scout Mountain Member. It is composed of metabasalts and volcanic breccias demonstrating chemistry consistent with intraplate, rift-related volcanism (Harper and Link, 1986). The Bannock Volcanic Member is broadly correlative along the Cordillera with other presumed synrift volcanics such as the Irene and Leola volcanics in NE Washington and the (ca. 685 Ma) Golden Cup volcanics in central Idaho (Lund et al., 2003). The Scout Mountain Member contains mostly siliciclastic rocks, including two glaciogenic diamictites separated by an apparently gradational interval of siltstone to cobble conglomerate (Crittenden et al., 1983; Link, 1983). The diamictites are considered to represent stades within a single ice age (Crittenden et al., 1983). Fanning and Link (2004) report three U-Pb sensitive high-resolution ion microprobe (SHRIMP) zircon ages from the Scout Mountain Member. A rhyolite clast within the upper Scout Mountain Member diamictite from Portneuf Narrows has been dated at 717 ± 4 Ma, and an epiclastic crystal tuff within the lower Scout Mountain Member on Oxford Mountain, ~60 mi (95 km) south of Pocatello, is 709 ± 5 Ma (Fig. 1). Thus, the upper glaciogenic diamictite, with its pink laminated dolomite cap carbonate (with negative carbon-isotope values) is younger than ca. 709 Ma by correlation. The dolomite cap carbonate is overlain by an erosional sequence boundary (boundary 1 on Fig. 3). In a scoured contact above an upward fining sandy succession is a 667 ± 5 Ma reworked tuff. This tuff is gradationally overlain by an upward fining succession and a second, cap-like carbonate also with negative carbon-isotope values (Lorentz et al., 2004). The (informal) upper member of the Pocatello Formation consists of >600 m of laminated argillite and/or shale with subordinate siltstone and quartzite (e.g., Crittenden et al., 1971; Link, 1983). The member becomes more silty upward and is overlain gradationally by siltstone to limestone cycles of the Blackrock Canyon Limestone. Thin limestone beds are present near the base of the upper member on the east side of Scout Mountain. The lack of cross stratification would suggest deposition below storm wave base. It is postulated that the upper member records a post-Sturtian global sea level rise event (Christie-Blick and Levy, 1989; Link and Smith, 1992).

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least five siliciclastic to carbonate cycles. The lower parts of these cycles contain substorm wave base, thin bedded siltstone, and fine-grained sandstone. The upper, carbonate parts of the cycles contain upward-shallowing microbial and oolitic carbonate beds. Hummocky cross bedding, internal erosion surfaces, and exposure surfaces are present in the upper, shallowest, parts of the cycles. Brigham Group The Brigham Group gradationally overlies the Blackrock Canyon Limestone (Figs. 1 and 3) (Crittenden et al., 1971; Link et al., 1987). It is up to 4000 m thick and consists of texturally and mineralogically mature strata, dominantly quartz sandstone, deposited in fluvial and shallow marine conditions (Link et al., 1993). It is composed of six formations: the Papoose Creek Formation, the Caddy Canyon Quartzite, the Inkom Formation,

Blackrock Canyon Limestone The Blackrock Canyon Limestone is positioned above the fining-upward Pocatello Formation and below the coarseningupward Brigham Group. It constitutes the third carbonate-bearing interval in the Portneuf Narrows succession (Figs. 1 and 3). The Blackrock Canyon Limestone, in its 530 ft (160 m) type section (Ludlum, 1942), is examined at Stop 2 (Fig. 5). Here it contains at

Figure 4. Stratigraphic columns of the Pocatello Formation in the Portneuf Narrows area. From Link (1987).

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Figure 5. Stratigraphic column of the Blackrock Canyon Limestone, near Pocatello, Idaho, with corresponding δ13C data (‰ variation referenced to the Peedee belemnite [PDB] standard). Carbonate intervals are numbered 1 (bottom) through 5 (top), as discussed in the text. Abbreviations: sil—siliciclastic; m—mudstone; w—wackestone; p—packstone; g—grainstone. Numerical data are shown in Corsetti et al. (2006).

Pocatello Formation and Overlying Strata, Southeastern Idaho the Mutual Formation, the Camelback Mountain Quartzite, and the Gibson Jack Formation (the latter of which is not plotted on Figs. 1 and 3). The Papoose Creek Formation contains distinctive irregularly banded very-fine grained quartzite and siltite, with abundant and distinctive subaqueous shrinkage cracks. The Caddy Canyon Quartzite is dominantly sandstone, marine at the base and fluvial in the upper part. In the middle of the Caddy Canyon Quartzite is a thin dolostone interval, which constitutes the fourth and final carbonate interval in the Pocatello area succession (Smith et al., 1994). Of note, deep incised valleys are reported in the upper Caddy Canyon Quartzite (denoted as sequence boundary 2 of Figs. 1 and 3). These have been considered to represent Marinoan-age glacio-eustatic drawdown (Levy et al., 1994). Incised valleys display tens of meters of erosional relief and are filled with fluvial sandstone and conglomerate. The Inkom Formation contains subwave base and turbiditic green phyllite and argillite with medial siltstones. The upper contact of the Inkom Formation is a sequence boundary of widespread significance (denoted as sequence boundary 3 on Figs. 1 and 3). The overlying Mutual Formation contains braided fluvial feldspathic and conglomeritic quartzites. Extrusive volcanics in correlative units (Brown’s Hole Formation) in Utah have been dated at 580 ± 7 Ma (Ar-Ar on hornblende from an alkali trachyte, Crittenden and Wallace, 1973; recalculated by Christie-Blick and Levy, 1989). The contact between the Mutual Formation and the Camelback Mountain Quartzite is sequence boundary 4 on Fig. 1 (e.g., Link et al., 1987, 1993). The Precambrian-Cambrian boundary is thought to reside in the Camelback Mountain Quartzite based on the presence of Cambrian trace fossils, but the precise position of the boundary is not known. The Camelback Mountain Quartzite is dominantly marine sandstone, whereas the overlying Gibson Jack Formation is marine siltstone and shale with sandy interbeds.

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bed in the section), directly overlying the upper glacial diamictite (<709 Ma) north of Portneuf Narrows. The unit is thus a classic Neoproterozoic “cap carbonate.” Lorentz et al. (2004) provided a detailed sedimentological study and δ13C profile (consistently negative δ13C values, ~−5‰) for the second (<667 Ma) carbonate interval, at the top of the Scout Mountain Member. This carbonate has “cap-like” features, including aragonite fans that may have formed on the sea floor, but it does not rest on a diamictite, and is part of a generally upward-deepening marine section that culminates in the laminated shale of the upper member Pocatello Formation. The Blackrock Canyon Limestone resides in the critical crossover (−2‰ to +1‰) from consistently negative δ13C values to consistently positive δ13C values. The crossover is located between siliciclastic to carbonate cycles 3 and 4 of the Blackrock Canyon Limestone (Fig. 5). As reported by Corsetti et al. (2006), the samples generally demonstrate significant recrystallization upon thin section analysis (although abundant primary allochems were apparent when Folk’s “white card” technique was used [Folk, 1987]). The δ18O of some of the samples are considered outside the range of well-preserved samples (e.g., Kaufman and Knoll, 1995). The δ13C is less susceptible to postdepositional alteration of this kind (e.g., Banner and Hanson, 1990). However, given the demonstrable postdepositional alteration of the δ18O values, the Blackrock Canyon data should be treated with caution when compared to other better-preserved Neoproterozoic sections. The fourth carbonate interval in the Pocatello area section, in the Caddy Canyon Quartzite, contains markedly positive δ13C values. Some of these reach +9‰ (Fig. 3), suggesting alteration, although Smith et al. (1994) interpreted them as primary values. In summary, the late Neoproterozoic strata in the Portneuf Narrows area record the change from mildly negative δ13C values above Sturtian age diamictites to highly positive values in overlying (sub-Marinoan?) strata (Figs. 1 and 3) (Lorentz et al., 2004). The data from the Pocatello succession suggests that the transition occurred after 667 Ma but before 580 Ma.

δ13C CHEMOSTRATIGRAPHY FOR THE POCATELLO SUCCESSION

Discussion

The Pocatello succession is not rich in carbonates, but it does contain an important set of radiometric dates. Shown in Figure 3 is a traditional correlation (following Miller, 1985; Link et al., 1993) of the Pocatello succession with the section in Death Valley. This correlation was questioned by Prave (1999), and until precise ages from the Death Valley succession are obtained, the issue remains controversial. Note that our correlation matches the cap dolomite above the upper diamictite of the Pocatello Formation with the Noonday Dolomite in Death Valley. Further, we correlate the −2 to +1‰ carbon isotope values of the Blackrock Canyon Limestone (and Brigham Group below the Inkom Formation) with the Johnnie Formation. Smith et al. (1994) report mildly negative δ13C values (−3‰) for the finely laminated pink carbonate unit (the first carbonate

The δ13C record between ca. 750 and 580 Ma is generally portrayed as highly enriched in 13C punctuated by negative δ13C excursions associated with Neoproterozoic glacial intervals (summarized in Melezhik et al., 2001). Hayes et al. (1999) show carbonates generally between ca. 700 and ca. 600 Ma as containing positive δ13C values. Some authors report a “geochemical divide” between the older Sturtian glacial interval and the younger Marinoan glacial interval, where δ13C values are ≥8‰ in between the two great ice ages (Kaufman et al., 1997; Knoll, 2000). Others have questioned the concept of the “geochemical divide” (e.g., Kennedy et al., 1998). Further, the Marinoan (<635 Ma) glaciation is now known to be distinct from the ca. 580 Ma Gaskiers-Varanger glaciations (Bowring et al., 2003; Hoffmann et al., 2004; Zhang et al., 2005; Condon et al., 2005).

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Integration of the available radiometric dates with the δ13C profile dictate that the duration of the post-Sturtian change from negative to positive δ13C values is much shorter (667–635 Ma; Fanning and Link, 2004; Zhou et al.,. 2004), at least in SE Idaho strata, than had been previously thought. Another possibility is that the caplike carbonate of the upper Pocatello Formation represents its own glaciation, one that has not been found elsewhere (Lorentz et al., 2004). Or, perhaps the negative isotope values and aragonite fans were formed in other conditions than during a postglacial alkalinity event (i.e., methane seeps; Jiang et al., 2003). DESCRIPTION OF STRATA EXAMINED AT STOP 1, NORTH OF PORTNEUF NARROWS Park in the flat area and climb the ridge to the northwest. Figure 2 shows the geology of the field trip stop area. The strata are structurally overturned, such that the stratigraphically highest rocks are at the top of the hill. Figure 4 shows the composite stratigraphy of the Pocatello Formation in the Portneuf Narrows area. The UTM coordinates of the gulch with the best exposures of the section from the diamictite up to the upper member are 12T 0389424, 4740031. The stratigraphically lowest rocks exposed are the upper diamictite of the Scout Mountain Member. These are matrix-supported clast-poor, unbedded diamictites, with cobble to boulder sized clasts (up to 1 m in diameter) of extrabasinal quartzite, granite, and gneiss, plus intrabasinal rhyolite, plagioclase-porphyritic basalt, and chloritic greenstone likely derived from erosion of the underlying Bannock Volcanic Member. Glacially striated clasts are present, though rare. Quartzite clasts are the best rock type to preserve the striations. An ~1-m-thick, finely laminated pink dolostone (carbonate number one) lies in abrupt unconformable contact with the uppermost diamictite (C. Dehler, 2005, personal commun.). An erosional sequence boundary (e.g., Link et al., 1993; sequence boundary 1 on Figs. 1 and 3) incises the dolostone, leaving behind dolostone breccia at some localities. The dolostone breccia is overlain by an ~100m-thick succession that fines upward from poorly sorted, mediumto coarse-grained arkosic sandstone containing planar and trough cross-stratification and dish structures to a succession of siltstone and lenticular sandstone containing asymmetric and climbing ripples, load casts, and sand lenses. Fanning and Link (2004) dated a reworked fluvial fallout tuff near the top of this succession at 667 ± 5 Ma (U-zircon) (Fig. 1). The tuff is the upper few cm of a 20 cm sandstone bed that lies on a scoured fluvial erosional surface. A 5 m limestone at the top of the Scout Mountain Member constitutes the second carbonate-bearing interval in the succession and is “caplike” in that it contains unusual cm-scale seafloor fans and negative δ13C values (Smith et al., 1994; Lorentz et al., 2004); however, it cannot be matched with any known glaciation. STRATA EXAMINED AT STOP 2, BLACKROCK CANYON We will examine the Blackrock Canyon Limestone, type locality (12T 0391615; 4740233. Inkom, Idaho, 7 1/2′ quadrangle).

Park on the east side of Blackrock Canyon Road and walk up the cyclic section directly to the east; a measured section is shown in Figure 5. It is an ~600 ft (183 m) climb to the top of the hill. Eaststriking cross faults offset the limestone beds. Despite its name, siltstones and sandstones dominate the Blackrock Canyon Limestone. Carbonates comprise only ~25% of the type section. The carbonate units thin southward along strike to only a few meters ~20 km to the south on the east face of Scout Mountain. The Blackrock Canyon Limestone in Blackrock Canyon contains siliciclastic and carbonate rocks arranged in five cycles (Fig. 5). Each cycle is composed of a siliciclastic lower part and a carbonate upper part. The lowest three cycles and the upper cycle are <20 m thick. The fourth cycle contains 90 m of siltstone below its upper carbonate bed. The lowest three cycles are the best exposed, but the upper (fifth) carbonate bed is by far the thickest (>65 m; Trimble, 1976). The siltstones of the siliciclastic part of the cycles are identical to those in the overlying Papoose Creek Formation. They contain small-scale cross beds and unusual subaqueous shrinkage cracks in thinly interbedded, parallellaminated siltstone and fine-grained sandstone (Crittenden et al., 1971). The carbonate units are predominantly laminated mudstones to packstones in the lower three cycles and cross-bedded oolitic grainstones in the upper two cycles. The lower carbonate units also contain poorly defined stromatolitic and/or thrombolitic bioherms on the order of a few meters in width and height. Hummocky cross bedding, erosion surfaces, and dolomitized exposure surfaces are present in the upper parts of the cycles. Each carbonate half-cycle represents shallower deposition than the preceding cycle (i.e., they forestep). We thus interpret the Blackrock Canyon Limestone cycles to represent initiation of highstand deposition. The parasequences (siliciclastic to carbonate cycles) thicken northward from Scout Mountain to Blackrock Canyon. This may reflect a tectonically controlled south-facing hinge on the flanks of the former Bannock Volcanic Member eruptive center. The thickest carbonates may have accumulated where subsidence was most rapid through the photic zone. ACKNOWLEDGMENTS The sensitive high-resolution ion microprobe geochronology for these rocks was supported by National Science Foundation EAR 0125756. Some of the text for this paper is modified from Corsetti et al. (2006). Editorial comments by Carol Dehler were very useful. REFERENCES CITED Banner, J.L., and Hanson, G.N., 1990, Calculation of simultaneous isotopic and trace element variations during water-rock interaction with applications to carbonate diagenesis: Geochimica et Cosmochimica Acta, v. 54, p. 3123–3137, doi: 10.1016/0016-7037(90)90128-8. Bowring, S.A., Myrow, P.M., Landing, E., Ramezani, J., Condon, D., and Hoffman, K.H., 2003, Geochronological constraints on Neoproterozoic glaciations and the rise of Metazoans: Geological Society of America Abstracts with Programs, v. 35, no. 6, p. 516.

Pocatello Formation and Overlying Strata, Southeastern Idaho Christie-Blick, N., and Levy, M., 1989, Stratigraphic and tectonic framework of upper Proterozoic and Cambrian rocks in the Western United States, in Christie-Blick, N., Levy, M., Mount, J.F., Signor, P.W., and Link, P.K., eds., Late Proterozoic and Cambrian tectonic, sedimentation, and record of metazoan radiation in the western United States, Field Trip Guidebook T331: Washington, D.C., American Geophysical Union, p. 113. Condon, D., Zhu, M., Bowring, S., Wang, W., Yang, A., and Jin, Y., 2005, U-Pb ages from the Neoproterozoic Doushantuo Formation, China: Science, v. 308, p. 95–98, doi: 10.1126/science.1107765. Corsetti, F.A., and Kaufman, A.J., 2003, Stratigraphic investigations of carbon isotope anomalies and Neoproterozoic ice ages in Death Valley, California: Geological Society of America Bulletin, v. 115, p. 916–932, doi: 10.1130/B25066.1. Corsetti, F.A., Link, P.K., and Lorentz, N.J., 2006, δ13C chemostratigraphy of the Neoproterozoic succession near Pocatello, Idaho, in Link, P.K., and Lewis, R.S., eds., Proterozoic Geology of western North America and Siberia: Society of Economic Paleontologists and Mineralogists (SEPM) Special Publication (in press). Crittenden, M.D., Jr., and Wallace, C.A., 1973, Possible equivalents of the Belt Supergroup in Utah, in Belt Symposium, v. 1: Moscow, University of Idaho, Idaho Bureau of Mines and Geology, p. 116–138. Crittenden, M.D., Jr., Schaeffer, F.E., Trimble, D.E., and Woodward, L.A., 1971, Nomenclature and correlation of some upper Precambrian and basal Cambrian sequences in western Utah and southeastern Idaho: Geological Society of America Bulletin, v. 82, p. 581–602. Crittenden, M.D., Jr., Christie-Blick, N., and Link, P.K., 1983, Evidence for two pulses of glaciation during the late Proterozoic in northern Utah and southeastern Idaho: Geological Society of America Bulletin, v. 94, p. 437–450, doi: 10.1130/0016-7606(1983)94<437:EFTPOG>2.0.CO;2. Fanning, C.M., and Link, P.K., 2004, U-Pb SHRIMP ages of Neoproterozoic (Sturtian) glaciogenic Pocatello Formation, southeastern Idaho: Geology, v. 32, no. 10, p. 881–884, doi: 10.1130/G20609.1. Folk, R.L., 1987, Detection of organic matter in thin-sections of carbonate rocks using a white card: Sedimentary Geology, v. 54, p. 193–200, doi: 10.1016/0037-0738(87)90022-4. Harper, G.D., and Link, P.K., 1986, Geochemistry of upper Proterozoic riftrelated volcanics, northern Utah and southeastern Idaho: Geology, v. 14, p. 864–867, doi: 10.1130/0091-7613(1986)14<864:GOUPRV>2.0.CO;2. Hayes, J.M., Strauss, H., and Kaufman, A.J., 1999, The abundance of 13C in marine organic matter and isotopic fractionation in the global biogeochemical cycle of carbon during the past 800 Ma: Chemical Geology, v. 161, p. 103–125, doi: 10.1016/S0009-2541(99)00083-2. Hoffmann, K.-H., Condon, D.J., Bowring, S.A., and Crowley, J.L., 2004, U-Pb zircon date from the Neoproterozoic Ghaub Formation, Namibia—Constraints on Marinoan glaciation: Geology, v. 32, p. 817–820, doi: 10.1130/G20519.1. Jiang, G., Kennedy, M.J., and Christie-Blick, N., 2003, Stable isotopic evidence for methane seeps in Neoproterozoic postglacial cap carbonates: Nature, v. 426, p. 822–826, doi: 10.1038/nature02201. Kaufman, A.J., and Knoll, A.H., 1995, Neoproterozoic variations in the C-isotopic composition of seawater; stratigraphic and biogeochemical implications: Precambrian Research, v. 73, p. 27–49, doi: 10.1016/0301-9268(94)00070-8. Kaufman, A.J., Knoll, A.H., and Narbonne, G., M., 1997, Isotopes, ice ages, and terminal Proterozoic earth history: Proceedings of the National Academy of Sciences (USA), v. 94, p. 6600–6605. Kennedy, M.J., Runnegar, B., Prave, A.R., Hoffmann, K.H., and Arthur, M.A., 1998, Two or four Neoproterozoic glaciations?: Geology, v. 26, p. 1059– 1063, doi: 10.1130/0091-7613(1998)026<1059:TOFNG>2.3.CO;2. Knoll, A.H., 2000, Learning to tell Neoproterozoic time: Precambrian Research, v. 100, p. 3–20, doi: 10.1016/S0301-9268(99)00067-4. Levy, M., Christie-Blick, N., and Link, P.K., 1994, Neoproterozoic incised valleys of the eastern Great Basin, Utah and Idaho; fluvial response to changes in depositional base level, in Dalrymple, R.W., Boyd, R., and Zaitlin, B.A., eds., Incised-valley systems; origin and sedimentary sequences: Society of Economic Paleontologists and Mineralogists (SEPM) Special Publication 51, p. 369–382. Link, P.K., 1983, Glacial and tectonically influenced sedimentation in the upper Proterozoic Pocatello Formation, southeastern Idaho, in Miller, D.M., Todd, V.R., and Howard, K.A., eds., Tectonic and stratigraphic studies in the eastern Great Basin: Geological Society of America Memoir 157, p. 165–181. Link, P.K., 1987, The Late Proterozoic Pocatello Formation; a record of continental rifting and glacial marine sedimentation, Portneuf Narrows, southeastern Idaho, in Beus, S.S., ed., Centennial Field Guide Volume 2: Rocky Mountain Section, Geological Society of America, p. 139–142.

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Link, P.K., and Smith, L.H., 1992, Late Proterozoic and Early Cambrian stratigraphy, paleobiology, and tectonics; northern Utah and southeastern Idaho: Field guide to geologic excursions in Utah and adjacent areas of Nevada, Idaho, and Wyoming, v. 92-3, p. 461–481. Link, P.K., Jansen, S.T., Halimdihardja, P., Lande, A.C., and Zahn, P.D., 1987, Stratigraphy of the Brigham Group (Late Proterozoic-Cambrian), Bannock, Portneuf, and Bear River Ranges, southeastern Idaho, in Miller, W.R., ed., The thrust belt revisited: 38th Annual Wyoming Geological Association Guidebook, p. 133–148. Link, P.K., Christie-Blick, N., Devlin, W.J., Elston, D.P., Horodyski, R.J., Levy, M., Miller, J.M.J., Pearson, R.C., Prave, A., Stewart, J.H., Winston, D., Wright, L.A., and Wrucke, C.T., 1993, Middle and Late Proterozoic stratified rocks of the western U.S. Cordillera, Colorado Plateau, and Basin and Range Province, in Reed, J.A. et al., eds., Precambrian; conterminous U.S.: Boulder, Colorado, The Geology of North America (DNAG), v. C-2, Geological Society of America, p. 463–595. Lorentz, N.J., Corsetti, F.A., and Link, P.K., 2004, Seafloor precipitates and C-isotope stratigraphy from the Neoproterozoic Scout Mountain Member of the Pocatello Formation, southeast Idaho: implications for Neoproterozoic earth system behavior: Precambrian Research, v. 130, p. 57–70, doi: 10.1016/j.precamres.2003.10.017. Ludlum, J.C., 1942, Pre-Cambrian formations at Pocatello, Idaho: Journal of Geology, v. 50, p. 85–95. Lund, K., Aleinikoff, J.N., Evans, K.V., and Fanning, C.M., 2003, SHRIMP U-Pb geochronology of Neoproterozoic Windermere Supergroup, central Idaho; implications for rifting of western Laurentia and synchroneity of Sturtian glacial deposits: Geological Society of America Bulletin, v. 115, p. 349–372, doi: 10.1130/0016-7606(2003)115<0349: SUPGON>2.0.CO;2. Melezhik, V.A., Gorokhov, I.M.K.A.B., and Fallick, A.E., 2001, Chemostratigraphy of Neoproterozoic carbonates; implications for “blind dating”: Terra Nova, v. 13, p. 1–11, doi: 10.1046/j.1365-3121.2001.00318.x. Miller, J.M.G., 1985, Glacial and syntectonic sedimentation; the upper Proterozoic Kingston Peak Formation, southern Panamint Range, eastern California: Geological Society of America Bulletin, v. 96, p. 1537–1553, doi: 10.1130/0016-7606(1985)96<1537:GASSTU>2.0.CO;2. Prave, A.R., 1999, Two diamictites, two cap carbonates, two δ13C excursions, two rifts; the Neoproterozoic Kingston Peak Formation, Death Valley, California: Geology, v. 27, p. 339–342, doi: 10.1130/0091-7613(1999)027<0339: TDTCCT>2.3.CO;2. Smith, L.H., Kaufman, A.J., Knoll, A.H., and Link, P.K., 1994, Chemostratigraphy of predominantly siliciclastic Neoproterozoic successions; a case study of the Pocatello Formation and lower Brigham Group, Idaho, USA: Geological Magazine, v. 131, p. 301–314. Stewart, J.H., 1970, Upper Precambrian and Lower Cambrian strata in the southern Great Basin, California and Nevada: U.S. Geological Survey Professional Paper 620, 206 p. Stewart, J.H., 1982, Regional relations of Proterozoic Z and Lower Cambrian rocks in the western United States and northern Mexico, in Cooper, J.D., Troxel, B.W., and Wright, L.A., eds., Geology of selected areas in the San Bernardino Mountains, western Mojave Desert, and southern Great Basin, California: Geological Society of America Cordilleran Section Volume and Guidebook: Shoshone, California, Death Valley Publishing, p. 171–180. Stewart, J.H., and Poole, F.G., 1974, Lower Paleozoic and uppermost Precambrian Cordilleran Miogeocline, Great Basin, western United States, Tectonics and Sedimentation: Society of Economic Paleontologists and Mineralogists (SEPM) Special Publication 22, p. 28–57. Stewart, J.H., and Suczek, C.A., 1977, Cambrian and latest Precambrian paleogeography and tectonics in the western United States, in Stewart, J.H., Stevens, C.H., and Fritsche, A.E., eds., Paleozoic paleogeography of the western United States: Society of Economic Paleontologists and Mineralogists (SEPM) Pacific Section Book 7, p. 1–17. Trimble, D.E., 1976, Geology of the Michaud and Pocatello Quadrangles, Bannock and Power Counties, Idaho: U.S. Geological Survey Bulletin 1400, 88 p. Zhang, S., Jiang, G., Zhang, J., Song, B., Kennedy, M.J., and Christie-Blick, N., 2005, U-Pb sensitive high-resolution ion microprobe ages from the Doushantuo Formation in south China: Constraints on late Neoproterozoic glaciations: Geology, v. 33, p. 473–476, doi: 10.1130/G21418.1. Zhou, C., Tucker, R.D., Xiao, S., Peng, Z., Yuan, X., and Chen, Z., 2004, New constraints on the ages of Neoproterozoic glaciations in south China: Geology, v. 32, p. 437–440, doi: 10.1130/G20286.1. Printed in the USA

Geological Society of America Field Guide 6 2005

Anatomy of reservoir-scale normal faults in central Utah: Stratigraphic controls and implications for fault zone evolution and fluid flow Peter Vrolijk Rod Myers Michael L. Sweet ExxonMobil Upstream Research Company, P.O. Box 2189, Houston, Texas 77252-2189, USA Zoe K. Shipton Center for Geosciences, University of Glasgow, Glasgow G12 8QQ, Scotland, UK Ben Dockrill Department of Geology, Trinity College, Dublin, Dublin 2, Ireland James P. Evans Jason Heath† Anthony P. Williams‡ Department of Geology, Utah State University, Logan, Utah 84322-4505, USA ABSTRACT Analysis of fault zone structure and composition of two intermediate-displacement faults in the Colorado Plateau reveal how fault structure varies as a function of lithology, and how faults impact fluid flow. The Little Grand Wash fault cuts Jurassic Summerville through Cretaceous Mancos Shale rocks, and consists of a complex set of interweaving fault strands. Fault relays are developed where sandstone and shale are cut by the fault in roughly equal amounts. Ancient and modern fluid flow is documented by the presence of travertine and tufa deposits, an oil seep, and CO2 gas seeps. Analysis of the water, travertine, and gas composition indicate that the CO2 emanates from a reservoir 1.5–2 km deep, and charges a shallow aquifer. Cross-fault flow is inhibited, and the gas and water flows in the footwall damage zone of the fault. Analysis of the Bighole fault in the San Rafael Swell shows how fault structure varies with displacement. The fault is exposed entirely in the aeolian Jurassic Navajo Sandstone, and consists of a dense fault core interpreted to be a densely packed set of deformation bands bounded by a narrow slip surface. The fault zone consists of conjugate deformation band sets in the hanging wall and footwall of the fault, and the fault core thickness does not vary significantly with net slip. Keywords: normal faults, fluid flow, fault mechanisms, clastic sedimentary rocks. ‡

Present address: Anadarko Petroleum Corporation, 1201 Lake Robbins Drive, The Woodlands, Texas 77380, USA Present address: Department of Earth and Environmental Sciences, New Mexico Institute of Mining and Technology, Socorro, New Mexico 87801, USA

†

Vrolijk, P., Myers, R., Sweet, M.L., Shipton, Z.K., Dockrill, B., Evans, J.P., Heath, J., and Williams, A.P., 2005, Anatomy of reservoir-scale normal faults in central Utah: Stratigraphic controls and implications for fault zone evolution and fluid flow, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 261–282, doi: 10.1130/2005.fld006(13). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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262 INTRODUCTION

fault zones and their impact on fluid flow, and we visit two sites that illustrate past and present flow and fault structure. This field trip focuses on field sites that show how the composition of fault zones (principal displacement surface, fault core, damage zone) varies depending on host rock lithology, how fault zones develop with increasing slip, and how fault zone lithologies and structure vary with the amount of slip. We also examine the influence of stratigraphic stacking patterns on fault nucleation and growth in a heterogeneous clastic sequence, and how faults affect CO2 and groundwater flow. The exposures examined in this field trip are field analogs for important faulted oil and gas reservoirs (such as the North Sea) and are also exhumed analogs to candidate sites for geologic sequestration of CO2.

Reservoir-scale faults are common structures that define the architecture of reservoirs. Faults in sedimentary sequences create complex heterogeneities that affect how fluid flow will occur in the reservoir and in fault zones. Clastic sedimentary rocks support water, gas, and oil, which flow predominantly through the permeable sandstone fraction. Where normal faults cut a sedimentary section, they alter the flow networks established by depositional and sedimentary processes. This is accomplished by: (1) disruption of the stratigraphic framework by fault offset; (2) destruction of sandstone pore fabrics by cataclasis; (3) separation of permeable sandstone beds by incorporating shale gouge into the fault zone; and (4) creation of permeable pathways along and around the fault surface that allow fluids to cross impermeable shales. We examine the evidence for fluid flow across and along faults at two locations in central Utah: the Little Grand Wash Fault near Green River, Utah, and the Big Hole fault on the northern end of the San Rafael Swell. We focus on key issues regarding

39°30’

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110° 30’

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STRUCTURAL AND TECTONIC SETTING The study area is in a region influenced by three tectonic provinces (Fig. 1): the Paradox Basin to the southeast, the Uinta Basin to the north, and San Rafael Swell to the west. The Paradox

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Unita Basin

BH

Uncompahgre uplift

San Rafael Swell Green River 39° LGW SW

Green River

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Moab fault

Figure 1. Regional geologic map of the field trip area. BH—Bighole fault; LGW—Little Grand Wash fault; SW— Salt Wash fault.

Moab La Sal Mtns.

Outline of the Paradox Basin Colorado River

38°

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30 km Major faults Rivers Boundaries of major structural elements

Anatomy of reservoir-scale normal faults in central Utah

Formation

Members

Map Thickness Lithology Symbol (Meters)

Upper Member Kmu 150+ Mancos Shale Ferron Ss Mbr. Km 15-35 Tununk Mbr. Kml 90-150 Dakota Sandstone Cedar Mountain Formation

Jurassic

Morrison Formation

Brushy Basin Mbr.

Kd Kcm

0-35 24-60

Jmb 90-135

Salt Wash Mbr. Jms 40-90 Summerville Formation Jmt 35-50 Curtis Formation Moab Mbr. Jctm 25-100 Entrada Sandstone Slick Rock Ss. Mbr. Jes Carmel Formation Dewey Bridge Mbr. Jcd

50-160 40-70

Navajo Sandstone

Jn

65-150

Kayenta Formation

Jk

75-120

Wingate Sandstone

Jw

55-160

Chinle Formation

TRc

60-95

Moenkopi Formation

TRm

0-150

Black Box Dolomite White Rim Ss Cutler Formation

Organ Rock Shale

Pcb 20-50 Pcw 90-150 0-90 Pco

Elephant Canyon Pce Formation

Honaker Trail Formation

Paradox Formation

Stratigrahy encountered on this field trip

Cretaceous

Age

We depart the Salt Lake City convention center on Wednesday afternoon for the 3.5–4 h drive to Green River, Utah, by taking I-15 south past Provo to Exit 261 for U.S. Hwy 6. Follow Hwy 6 over Soldier Summit, through Price, and to I-70. Take I70 east for 2 mi to Exit 158 for Green River, Utah (Fig. 3).

Triassic

The rocks that are the focus of this field trip are from the Jurassic and Cretaceous section of the central Colorado Plateau (Fig. 2). The Jurassic section in this part of Utah is dominated by nonmarine deposits. Arid conditions prevailed during the early Jurassic and middle Jurassic, witnessed by deposition of thickly bedded aeolian sandstones such as the Wingate, Navajo and Entrada Formations. In the mid- and late Jurassic the climate became more humid with deposition of fluvial and lacustrine units in the Morrison Formation. Marine incursions into this dominantly nonmarine succession are recorded by the Summerville, Curtis, and Carmel Formations. The Jurassic section is punctuated by a number of regionally

FIELD GUIDE

Permian

STRATIGRAPHIC SETTING

correlative unconformities (J-0 through J-5), mapped in the region by Pipiringos and O’Sullivan (1978). A major unconformity occurs at the top of the Morrison Formation and separates it from fluvial deposits of the Lower Cretaceous (Albian) Cedar Mountain Formation. Above the Cedar Mountain Formation is a major unconformity that encompasses the Middle Cretaceous and heralds the arrival of the Western Interior Seaway with marine rocks of the Upper Cretaceous (Cenomanian) Dakota Sandstone and Mancos Shale. These are the youngest rocks cut by faults visited in this trip.

Pennsylvanian

Basin is a Pennsylvanian intracratonic basin that accumulated a thick sequence of carbonate, clastic, and evaporite sediments. The basin extends as far north as the town of Green River and is regionally characterized by a series of northwest-trending folds and parallel faults. The Uinta Basin is to the north, with its southern boundary defined by the Book Cliffs escarpment. The strata gently dip north to northeast from this escarpment to the center of the basin. The San Rafael Swell is an arcuate, asymmetric, doubly plunging anticline that lies to the west of both basins. The anticline has a broad, gently dipping west limb and a steep east-dipping limb that forms the San Rafael monocline. The San Rafael Swell is a passive drape fold that developed above a reactivated basement reverse fault during Laramide contraction (50 Ma). Within the northernmost part of the Paradox Basin (Fig. 1), the principal structural features are a set of west-northwest–trending normal faults (Little Grand Wash fault, Salt Wash Graben and Ten Mile Graben), the northwest-trending Green River anticline and Courthouse syncline. The Green River anticline and Courthouse syncline form part of a series of northwest-trending folds in the northern Paradox Basin that have growth histories related to salt migration since the Permian (Doelling, 1988). The Green River anticline is a north- to northwest-trending open anticline that is cut by a set of normal faults. The axis of the anticline plunges shallowly to the northwest. The WNW-trending, 60–80° dipping Little Grand Wash and Salt Wash normal faults cut the northern end of the Paradox Basin. Timing of continued movement along these faults is poorly constrained. Pevear et al. (1997) and Williams (2005) present evidence for Early Tertiary and Quaternary slip on these faults. The faults cut the Mancos Shale consistent with substantial fault activity having occurred at least up to the Middle Cretaceous. The faults cut a north-plunging anticline, which could be related to salt movement in the Paradox Formation at depth. A basin-wide system of salt anticlines initiated when the salt was loaded by the Pennsylvanian and Permian clastic sediment shed off the Uncompahgre uplift to the northeast. Reactivation of the salt-related anticlines and faults occurred during Laramide (Eocene) contraction (Chan et al. 2000).

263

300-365

lPh

230+

lPp

760+

Figure 2. Stratigraphic section of the central Colorado Plateau area.

P. Vrolijk et al.

264

The entire first day is devoted to the Little Grand Wash fault, starting with a stratigraphic overview of the Jurassic Summerville and Morrison, and Cretaceous Mancos Formations. This records a transition from a fluvial/lacustrine to marine conditions. This discussion emphasizes the stratal and stratal packaging geometries throughout this succession. We examine the Little Grand Wash fault where it is exposed along a 2 km strike-section. Although the fault appears simple at a gross scale, complexity is developed at a range of scales, yielding both strike- and diprelays along the fault trace. During the field trip, we will discuss ideas for how the original stratigraphic assemblage influences the structural complexity, and finally speculate on how that structural evolution may influence cross-fault fluid flow. Lastly, the trip

examines dramatic evidence for along-fault fluid flow and CO2 migration, defined by large travertine mounds and pavements. This trip starts at the Holiday Inn Express in Green River, Utah. Road Log to Stop 1.1 Mileage Interval (Cumulative) Description 0.0

(0.0)

1.8

(1.8)

Proceed east on Main Street, out of town, over I-70, and to a T-intersection. Turn right and proceed up the hill.

Legend

A !(

Great Salt Lake

(!Logan !

Ogden !( !( !(

(! Salt Lake !

County Line Freeway Waterbody Main Elevation 13,000 feet Route for 2500 feet field trip

!( Uinta Mountains !(

!(

!( Provo

!(

B ! Price

!( !(

San !Rafael Swell Day 2

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Green River Day 1

!(

!( !(

!( St George

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Figure 3. Geographic setting of the field trip. (A) Shaded relief map of Utah, with major highways and the route from Salt Lake City to Green River, Utah, shown. (B) Details of Day 1 field trip, from the San Rafael Desert 30′ × 60′ quadrangle. (C) Details of Day 2 field trip, from the Huntington 30′ × 60′ quadrangle.

Anatomy of reservoir-scale normal faults in central Utah 0.4

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Near the crest of the hill, turn hard left onto a dirt road. You will immediately pass a patterned array of concrete pads left over from the former Utah Missile Range (mile 2.4). Continue up a gentle hill. Just beyond the crest of the hill, you pass the Buckhorn Conglomerate Member of the Cedar Mountain Formation, the first fluvial sandstone below the Mancos shale, which was deposited in the Cretaceous interior seaway (mile 3.2). Continue along this road across a flat plain surrounded by low outcrop hills. Bear right at a Y-intersection, take a hard right at a bunker for a former missile launching pad, and proceed to the tall radio tower. Base of steep road up to radio tower. Pending road conditions and car type, we may choose to either park at the base of this hill or proceed up to the radio tower. This is a steep, narrow road with some overhanging outcrops, so proceed with caution if driving.

Stop 1.1: Stratigraphic Overview At this stop, ~400 m of Lower Jurassic through Upper Cretaceous rocks are exposed in the footwall of the Little Grand Wash fault (Fig. 4). In the hanging wall of the fault only the upper part of this section is exposed. The oldest rocks exposed are near Crystal Geyser. They are part of the Middle Jurassic (Collovian) Summerville Formation. The Summerville Formation is composed of interbedded, brick-red, sandstones, siltstones and mudstones with gypsum nodules. Prominent coarsening- and thickening-up parasequences are characteristic of the Summerville Formation, which was deposited in a marine sabkha environment. A subtle, angular unconformity is observed at the top of the Summerville separating it from the overlying Tidwell member of the Morrison Formation. The lowermost member of the Upper Jurassic (Kimmeridgian–Tithonian) Morrison Formation is the Tidwell. It comprises ~20 m of interbedded, green, red, and purple mottled sandstone, siltstone, and mudstone. Most of the structures developed along the Little Grand Wash fault are in the upper two members of the Morrison Formation—the Salt Wash and the Brushy Basin. The Salt Wash Member of the Morrison Formation is ~100 m thick at this location. It is characterized by thickly bedded, cross-stratified sandstones and conglomerates interbedded with mottled mudstones and siltstones. The Salt Wash in this part of Utah is widely interpreted as the deposits of braided streams and their associated floodplain deposits (e.g., Peterson, 1984). These

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rivers flowed from a rising mountain belt to the west into a series of lacustrine basins in the vicinity of Grand Junction, Colorado (Turner and Peterson, 2004). Within the Salt Wash sandstones and conglomerates deposits are interpreted as the result of fluvial bars that are stacked into larger-scale sedimentary bodies called bar complexes. The abundance and continuity of conglomerate and sandstone bodies relative to the siltstone and mudstone varies both vertically and laterally depending on the stacking arrangement of these bars and bar complexes and the angle at which they intersect the fault plane (Fig. 4). For example between Sections 1 and 2 the Salt Wash and Brushy Basin Members are relatively sandy, but there is a significant reduction in sand content west and east of these sections (Fig. 4). In the vicinity of Section 4 note that the aspect ratio of the sand bodies increases relative to the west end of the outcrop. Paleocurrent measures show that in this part of the fault zone the fluvial bars are intersecting the fault plane obliquely, while around Section 1 the bars are more perpendicular to the fault plane. The top of the Salt Wash Member is marked by an unconformity (Demko et al., 2004) that is subtle here and a significant change in the net-to-gross of the section (Fig. 5). Overlying the Salt Wash Member is the Brushy Basin Member (Fig. 4). The Brushy Basin Member is distinguished from the Salt Wash Member by its much lower ratio of sandstone and conglomerate to mudstone. It is characterized by prominent red and white muddy siltstones, interpreted as paleosols, some of which contain abundant volcanic ash. In the footwall of the Little Grand Wash fault, it is ~130 m thick. The Brushy Basin Member fines upward and sandstone and mudstone-filled channels are found in this interval. A thin, laterally extensive conglomerate occurs near the top of the Brushy Basin in this area (Fig. 5). The depositional environment of the Brushy Basin Member is interpreted to have been meandering streams and floodplains with a higher water table than the underlying Salt Wash Member (Demko et al., 2004; Turner and Peterson, 2004). Dinosaur bones are common within some of the sandstones of the Brushy Basin Member near the Little Grand Wash fault. A major unconformity occurs at the top of the Salt Wash Member and marks the transition between Late Jurassic and Early Cretaceous. The Lower Cretaceous Cedar Mountain Formation was deposited in a fluvial environment. In the vicinity of the Little Grand Wash fault the Buckhorn Conglomerate Member is very prominently exposed. This conglomerate is ~20 m thick and makes an excellent marker bed, as it is continuous over the study area and caps many small hills. Capping the Cedar Mountain Formation is a major unconformity that encompasses the entire middle Cretaceous (Doelling, 2002). The Dakota Sandstone rests on this unconformity. It is thin and poorly exposed in the hanging wall and footwall of the Little Grand Wash fault. A characteristic feature of the Dakota Sandstone near the fault is the occurrence of a thin conglomerate with clasts of bivalve shells and chert. Overlying the Dakota are the thick, gray mudstones of the Mancos Shale.

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Anatomy of reservoir-scale normal faults in central Utah Mileage Interval (Cumulative) Description

Road Log to Stop 1.2 Proceed down the hill from parking spot below the radio tower.

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Pending road conditions and vehicle type, it is possible to drive along strike for ~2 mi to reach the main dirt road down to Crystal Geyser. This road takes you along a narrow, rutted road (4-wheel drive with reasonable clearance recommended) through the Lower Brushy Basin Member of the Morrison Formation. Continue back along the road we came in on, turning right at the paved road. T-intersection near entrance to I-70. Continue straight along the perimeter of the former Utah Missile Range. Turn right at intersection, following road to Crystal Geyser Optional Stop 1.2.

Stop 1.2: Cedar Mountain Formation—Buckhorn Conglomerate (optional) The Buckhorn Conglomerate Member is cut and well exposed in the road cut here. Note the grain size, trough crossbedding, bed thickness, and calcite cement expressed at this location. About 100 m in an up-section direction on the east side of the road, a mixed lithology of black-coated pebbles is exposed as a lag surface on top green and brown shales. These pebbles are exposed in outcrops in bluffs down near the river as a stratigraphic horizon. We interpret this surface as a sequence boundary–flooding surface correlative with the Dakota sandstone, which otherwise is absent from this area. The stop also allows a further examination of the lithologies and stratigraphic geometries present in the Brushy Basin Member of the Morrison Formation. From this vantage point, we see that the Brushy Basin Member consists of a large expanse of shale with thin, laterally discontinuous channelized sandstones. The thicker sandstones capping hills on the intermediate horizon are the Buckhorn Conglomerate Member. Note the layers of white shale exposed within the hills in the near distance. These shales are dominated by smectite and are interpreted as ash-fall deposits. The thickness of these beds changes throughout the field area. Part of this thickness change is due to bed erosion by younger intervals, but part is owed to reworking of the air-fall deposits through a network of fluvial channels. Note the gentle northern dips of the strata at this point. Road Log to Stop 1.3 Drive downhill and down stratigraphic section.

Little Grand Wash fault—easternmost exposure. Stop and park along the outside shoulder of a switchback in the road. Walk ~70 m up a steep, worn, single-file trail on the ridge of the Mancos Formation. Follow the tire tracks to one of the western promontories on this plateau.

Stop 1.3: Little Grand Wash Fault Overview Closer examination of the outcrops indicates that the fault is more complex than the regional picture suggests. Fault-bounded blocks with bedding dips to the east (along the fault strike) and to the south (toward the hanging wall). Fault relays are evident from the detailed map pattern (Figs. 6 and 7). Although strike relays are readily identified in surface exposures (Fig. 7) and are characterized by steep branchlines (fault intersection lines), there is also evidence for dip relays with shallow branchlines that we will examine later. However, the map pattern (Fig. 7) reveals that fault relays also exist without becoming breached, and both strike and dip relays exist without branchlines but with fault segment overlap. This vantage point also provides a good opportunity to view the three-dimensional stratigraphic complexity of the sandstone bodies in the Salt Wash Member. Compensational stacking of bars and bar complexes and rapid lateral changes in sandstone thickness is evident in these outcrops.

Stop 1.4b Stop 1.4a

Figure 6. Overview of east end of fault, with fault traces superimposed on high-resolution air photo draped onto the 30 m digital elevation model of the area. Image illustrates that fault zone consists of fault segments that intersect with vertical branchlines, creating fault relays, and segments that end in fault tips. Fault network defined by surface mapping, high-resolution differential Global Positioning System, and fault modeling in gOcad.

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The variability in the faulted stratigraphy, both in a vertical and lateral sense, and the change in the along-strike character of the fault zone has led us to hypothesize that the contrasting mechanical properties of the sandstone and shale lithologies and the three-dimensional distribution of those lithologies are responsible for the complexity of the fault network. Furthermore, we hypothesize that when a fault cuts through a dominantly sandstone or shale sequence, the fault is simple and few relays are formed. Where the sequence is heterolithic vertically and laterally, fault segments that initially nucleate in a number of sandstone beds (interpreted to fail in a brittle mode) will be misaligned and will coalesce through fault evolution into a series of fault relays. The remainder of the day is devoted to examining outcrops that were used to document the stratigraphic variability, the character and the distribution of fault relays, and to evaluating the relationship between the two.

the dip of each fault segment is less than the composite fault zone dip) (Fig. 10). Further evidence for this interpretation comes from comparing fault dips measured in outcrop (70–80°) to those from cores of this fault (60–65°), which intersect the fault as deep as 300 ft (91.5 m) below the surface.

Road Log to Stop 1.4

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(12.7)

Pull off on the right side of the road and walk 200 m NE to a small excavated fault exposure. Then walk W along the fault 100 m to a larger excavated fault exposure—be careful to walk around this outcrop; walking across it could cause it to collapse.

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(13.0)

Pull off on the right side of the road and walk 200 m east into the mouth of a small dry-wash feeding out from the cliffs.

Walking the wash toward the outcrop, we pass a large, steep hill of intensely veined Brushy Basin sandstones and shales capped by thick travertine deposits. This is a record of the CO2 migration along the Little Grand Wash fault and is the subject of later stops today. Past this tremendous outcrop, pause at the head of the sandy part of the wash to observe a brick-red, very fine-grained sandstone ledge. This is the first outcrop of the older Summerville Formation. The shales and thin sandstones in the hillside above us, up to the first thick channelized sandstone body, belong to the basal Tidwell Member of the Morrison Formation. The outcrop on the west side of the wash contains a segment of the Little Grand Wash fault, in this case filled with shale.

Stop 1.4 This stop consists of two outcrops: a and b (Fig. 6). At the first stop (1.4a), the mapped expression is simple, with most of the stratigraphic offset (Salt Wash on Mancos) developed on a single fault surface (Fig. 8). A second, smaller fault drops sandstones and shales of the lower Brushy Basin Member against the Salt Wash Member, but the main fault surface is developed as a simple fault (Fig. 8). In contrast, Stop 1.4b has a fault zone 15 m wide (Fig. 9). Deformation within this fault zone is variable, with steeply dipping fabrics, scaly clays, and closely spaced faults at the hanging wall and footwall cutoffs. In between we find locally continuous, thin sandstone beds broken into meter-sized fault blocks with subhorizontal bedding dips. We interpret most of the strain to have accumulated at the margins of this fault zone and the relatively undeformed center to be the last remnant of the fault sliver exposed to the east. This outcrop also allows us to examine evidence for dip relays. Looking back to the east, it can be seen that the rocks preserved in the fault relay dip toward the hanging wall. The bedding dips toward the hanging wall result when misaligned, overlapping fault segments develop in a releasing bend geometry (i.e.,

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Anatomy of reservoir-scale normal faults in central Utah Travel up the wash toward the west and onto the sandstone ridge along the south side. From a vantage point above the head of the wash, observe gentle north dips in sandstone beds of the basal Salt Wash Member on the north side of the wash, and moderate bedding dips to the east on the ridge. Moreover, continuous bedding dips are traced around the head of the wash and down the ridge, leading to the interpretation of a fault tip in an unbreached fault relay. The fluid-flow implications of this fault geometry are also clear from this location. Looking back to the north side of the wash, sandstone beds are separated by shales, limiting the amount of vertical flow in this aggregate section. Based on the limited lateral extent, both in cross section and map pattern, for these sandstone intervals, we expect numerous dead-end flow paths. However, this fault geometry introduces numerous short-circuits across the shale beds. The sandstone bed we are standing on touches several deeper sandstone intervals across the fault surface, allowing easy flow communication with adjacent stratigraphic levels. Looking across to the south side of this ridge we see the other fault segment of this relay and how the total displacement along the Little Grand Wash is distributed across a wide zone of fault segments and relays. There are substantially different expressions of the Little Grand Wash fault along the short segment we have examined. Other outcrops show variations on the principal observations we have made: (1) variations in fault zone thickness and complexity, defined by number of fault segments and variations in bedding dip within fault-bounded slivers; (2) strike relays, characterized by mappable fault tips and bedding dips along the strike of the fault; and (3) dip relays with bedding dips toward the hanging wall and with fault segments offset by tens of cm to m. The maximum stratigraphic offset (near the radio tower) across the Little Grand Wash fault is 195 m, thus the ratio of bed thickness to fault throw is 0–200, an uncommon range for such a large exposure. Many of the fault exposures we examine thus have fault throws that are only several multiples of bed thickness. This leads us to the speculation that bed thickness may play a role in the fault geometry evolution. To explore this idea further, a detailed two-dimensional stratigraphic framework on the footwall from the Salt Wash Member to the Cedar Mountain Formation (Fig. 10) divides the section into sandstone or shale lithologies and sums the sandstone fraction for various bin sizes (Fig. 11). This is posted along a fault map (Fig. 11), from which a first-order correspondence emerges between low net:gross (N:G; i.e., sandstone fraction) sections and simple fault development and intermediate N:G corresponding to multi-stranded fault intervals with bedding dips reflecting strike- and dip-relays. If this fault were more deeply exposed, we anticipate that the fault through the Summerville Formation would be simple. Fault relays and fault complexity arise from faults propagating through a section of mixed sandstone and shale lithologies in roughly comparable proportions and thus thickness. Moreover, sandstones and shales must have contrasting mechanical

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properties so that the applied stresses cause the sandstones to deform brittlely while the shales deform ductilely. The shales must also be sufficiently ductile that any stress perturbations arising from incipient faulting in sandstone beds is either absorbed in the shale interval or deflected away from the sandstone failure plane. This allows brittle failure in the adjacent sandstone bed to occur out of the plane of the initial failure, setting the stage for fault relays to develop with continued fault offset. The mechanical implications of the proposed fault evolution hypothesis are consistent with the burial history of these rocks. Based on analogy with similar and older rocks along the Moab fault to the southeast (Pevear et al., 1997) and the fault age determined by dating illites in the fault here (38 Ma), we infer that at least some of the fault history occurred at depths of ~2000 m. At these depths, the sandstones would be sufficiently consolidated to deform in a brittle mode, whereas the shales in this section, sometimes dominated by volcanic ash and smectite, would deform in a more ductile manner. The interpretation of the rocks in the outcrops and the interpretation of their structural history open an interesting and largely unexplored area of research into the mechanics of deformation of layered, mechanically heterogeneous rocks. There is an interesting interplay between the geometric distribution of the stratigraphic elements in the context of their evolving mechanical properties. Would the same fault pattern have developed through this section if faulting had occurred at Earth’s surface? The fluid flow implications of our fault evolution hypothesis are intriguing. Considering the difficulty in finding flow connections in the heterogeneous Salt Wash Member, separated both vertically and laterally by shales, faulting has improved flow connections. The heterogeneous distribution of sandstone and shale that hinders flow potentially promotes fault relays and associated bedding dips that establish new flow connections (if faulting occurs for the appropriate mechanical properties). If our faulting hypothesis can be substantiated, then we have the potential to better risk flow connections knowing only the stratigraphic distribution of sandstone and shale and the relative mechanical properties of those lithologies during faulting. Road Log to Stop 1.6 Continue farther west on the road toward the river. Mileage Interval (Cumulative) Description 0.2

(13.2)

0.3

(13.5)

Veer right at the Y-intersection of the road. Park at the southern end of the travertine terrace near the abandoned well. We will first visit the active travertine deposit associated with the abandoned well and then walk 200 m to the east to examine complex relationships between the fault and ancient travertines.

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Figure 10. Simplified fault map with schematic cross sections near Stop 1.4b. South-dipping beds are interpreted to represent a fault dip-relay. Two different scenarios for the position of the tiplines are described for section Y–Y′.

Figure 11. Upper figure is a stratigraphic section of footwall, expressed as N:G (net:gross; i.e., sandstone fraction) in 100 ft bins. Lower figure is an aligned fault trace map. Letters denote locations of correlations between the stratigraphic section and the trace map. (A) “Thick” fault zone correlates to regions of high N:G in the lower Brushy Basin and Salt Wash Members. (B) “Thin” fault zone (simple architecture) correlates to regions of lower N:G. (C) Prominent relay in lower Brushy Basin Member correlates to increased lower Brushy Basin Member N:G. (D) Shale gouge-prone relay in SW correlates to region of “layered” mechanical stratigraphy. (E) Abrupt thinning of fault zone correlates to abrupt N:G change in lower Brushy Basin Member (F). Buckhorn conglomerate Member (Cedar Mountain Formation) between fault strands occurs where isolated high N:G (stiffer unit) is embedded in shale rich section.

Anatomy of reservoir-scale normal faults in central Utah Stop 1.6 The Little Grand Wash and Salt Wash faults (~9 km to the south of Little Grand Wash fault) are the focus of a series of active CO2-charged springs (Fig. 12). The most dramatic of these springs is the Crystal Geyser system, which is located a on the eastern bank of the Green River ~50 m north of the Little Grand Wash fault zone. This cold-water geyser erupts to heights of up to 25 m at 4–12 h intervals and is powered by CO2-charged waters. The geyser began erupting when the Glen Ruby #1-X well was drilled in 1935 through a 21.5-m-thick travertine mound to the base of the Triassic section (Baer and Rigby, 1978). Three springs immediately north and northeast of the travertine deposit erupt periodically throughout the geyser eruption cycle and appear to be related to the same spring system. Evidence for active fluid flow is seen in several other locations along the fault. Small intermittent CO2 fluxes can be observed in the Green River, where a small line of bubbles follows the trace of the fault. Dry gas can be heard to seep from one location. Further to the east, an oil seep is located on the southernmost fault strand of the Little Grand Wash fault. A shallow pit

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contains fresh oil, indicating that there is active flow of petroleum to the surface. The outcrop close to this seep (Salt Wash Member of the Morrison Formation) contains patches of bitumen staining. Without a detailed diagenetic study it is not clear if the flow of hydrocarbons is related to the flow of CO2, but the close spatial association of the oil seep and the CO2 leaks suggests that similar pathways are being used by the hydrocarbons and the CO2. Travertine deposits are developed to various degrees around all the active springs. Well-developed mounds consist of a series of downstepping lobes that radiate out downslope of the springs, covered in terraces or rimstone, each with an undulating pool and a raised rim. The ochre color of the travertines indicates the presence of iron in the erupting waters. Attempts to core the active travertine were unsuccessful, as it appears that much of the mound is composed of organic material that has little strength. Less well-developed travertine mounds are formed round a number of the natural springs and wellbore leakage sites along the Salt Wash fault and we suggest that the degree of travertine development reflects the length of time that spring has been active. A detailed petrological and isotopic analysis of the travertines is summarized in Dockrill et al. (2006). Heath (2004)

Figure 12. Local geological map of the distribution of active springs, travertine and reduction zones along (A) the Little Grand Wash and (B) the Salt Wash faults. After Dockrill et al. (2006); Williams (2005) and Doelling (2001). Note that geologic formations have been grouped to simplify the map.

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examined the geochemistry of the gas-water-carbonate system along the Little Grand Wash and Salt Wash faults. We briefly summarize some of that study’s observations regarding the carbonate deposits and geochemistry of the system. Structure of the Travertine Mounds A series of ancient travertine mounds parallel the trace of the Little Grand Wash fault. All the travertines along both faults are localized either on the northernmost fault trace or in the footwall of these fault zones. The ancient mounds are up to 4 m thick and occur at elevations up to 30 m above the present level of the actively forming deposit, forming a resistant cap on top of a series of buttes. Dockrill et al. (2006) identifies five distinct lithofacies in the ancient travertines: (1) altered host rock, (2) white banded veins, (3) brown banded veins, (4) conglomerate, and (5) rice paddy unit (Fig. 13). The host rock underlying the ancient travertine deposits is crosscut by carbonate veins and altered along sporadic beds and surrounding subvertical fractures. In sand-dominated lithologies, the veins are 4–15 cm thick and usually alter the surrounding rock to produce a reduction halo 1–5 cm wide. In mud-dominated lithologies, the thick veins have mainly been substituted

for a network of thin boxwork veins (5–20 mm thick) that can obliterate the host rock fabric (Fig. 13). Larger areas of reduced sandstone are visible in the Curtis and Entrada sandstones. Dockrill et al. (2006) and Heath (2004) show that the geochemistry of the carbonates and altered host rock and their stable isotope composition is consistent with them being formed from similar waters to those effusing from the active springs. Hydrogeochemical Investigations Water samples from the springs in the area had in situ temperatures <18 °C. The waters are slightly acidic and very saline, with pH values ranging from 6.07 to 6.55 and total dissolved solids (TDS) values ranging from 13,848 to 21,228 mg/L. All of the waters are in the sodium-chloride chemical facies. The waters are supersaturated with respect to calcite, aragonite, dolomite, and hematite, and are undersaturated with respect to anhydrite, gypsum, halite, and quartz. Examination of waters from springs on both sides of the faults, show that only the springs in the footwalls of the faults are enriched in CO2. Hydrogen and oxygen isotopic data (from Heath, 2004) from water samples show that the waters are sourced from shallow groundwater.

Figure 13. Ancient travertine deposit capping the top of a butte near Crystal geyser, facing west. (A) The uppermost porous “rice paddy” unit; pencil is 15 cm long. (B) Dendritic brown-banded vein cross-cutting a conglomerate; lens cap is 8 cm in diameter. (C) White banded vein; field of view is 15 cm high. (D) Underlying host rock (Summerville formation) with dense carbonate veins. (E) Grainsupported conglomerate; lens cap is 8 cm in diameter. All photos by Ben Dockrill.

Anatomy of reservoir-scale normal faults in central Utah The gases emanating from all the springs are CO2 rich (95.66% to 99.41%) with minor amounts of Ar, O2 and N2. The δ13C values of the CO2 gas phase show little spread, ranging from −6.42 to −6.76‰ with a standard deviation of 0.13‰. The R/Ra value of Crystal Geyser spring is 0.302 (R = 3He/4He ratio of a sample and Ra = 3He/4He of the atmosphere). Since some atmospheric entrainment may have occurred, the actual R/Ra value of the emanating gases may be lower than these values. Crustal gases typically have R/Ra values of ~0.02, whereas mantle derived gases typically are ~8 (Kennedy et al., 1997). Thus, Heath (2004) concluded that the CO2 has a crustal source. Possible processes that may have generated the CO2 gas in the field area include (Cappa and Rice, 1995): (1) mantle or magmatic emanations; (2) the degradation of organic matter; (3) diagenetic reactions involving clay and carbonate rocks; and (4) thermal decarbonation of carbonate rocks by metamorphic processes. On the basis of detailed carbon isotope fractionation studies, Heath (2004) concluded that clay-carbonate diagenetic reactions that occurring during deep burial of impure carbonate sedimentary rocks are a likely source for the bulk of the abundant CO2 gas in the region. Travertine at the Crystal Geyser is composed of a mixture of calcite and aragonite closer to the geyser vent while distally the travertine is composed almost entirely of calcite. In the ancient travertine, the white banded veins are composed entirely of aragonite while the brown banded veins contain a mixture of aragonite and goethite with minor amounts of calcite. The active Crystal Geyser travertine has δ13C values ranging from 5.2 to 9.5‰ and δ18O values ranging from 18.0 to 19.4‰. At the base of this deposit the normally dark purple-red Summerville Formation is bleached to a pale yellow for up to several meters into the footwall. This alteration is localized along the subsidiary faults and within certain beds. The altered host rock exhibits δ13C and δ18O values between 2.7 and 3.8‰ and 19.9–22.7‰, respectively. Flow System Interpretation We construct a conceptual model of the regional groundwater flow in the upper 1.5 km of the basin (Fig. 14). Potentiometric surface data from groundwater wells show that regional groundwater flows from the northwest to the southeast (Hood and Patterson, 1984). Water temperature and stable isotope data for springs along both faults show that the CO2 is charging a reservoir ~300–500 m below the surface. All of the modern and ancient CO2 leakage points lie on the structural high where the north-plunging anticline is cut by the faults. Therefore, the faults are acting as flow barriers to southeast-directed CO2-charged groundwater flow, and CO2-charged groundwater is accumulating against the faults within the folded reservoir. Within the framework of this model for the geometry of the shallow CO2 storage system, the observations and data collected at the leaky faults can give us insight into each of the three components of the CO2 system, which are as follows: 1. CO2 reservoir. In our model, shallow groundwater reservoirs are charged from below by CO2 generated at depth. These

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shallow reservoirs are the high-porosity aeolian or fluvial Jurassic sandstones. The extent of reaction between the reservoir rocks and the CO2-charged fluid is unclear, but the continued effusion of the springs and geysers at the same locality through time shows that the high porosity must be maintained, and that the porosity is not being clogged by the products of diagenetic reactions. 2. Cap rock and fault seal. The topseal for upward movement of fluid from the shallow aquifer reservoirs is provided by shalerich Jurassic units. A lateral seal is provided by the fault rocks. The difference between the stable isotope signatures of veins in the hanging wall and footwall of the faults shows that the faults have acted as effective barriers for cross-fault flow at depth. The nature

Figure 14. Schematic block diagram cross section through the Little Grand Wash fault after Williams (2005) and Shipton et al. (2005), used to illustrate the conceptual reservoir model developed in these studies. The subsurface geology, specifically unit thicknesses, is constrained by wells. CO2-charged groundwater (small circles) is pooled in a north plunging anticlinal trap against the south-dipping fault. Although the fault zone geometry is much more complex than is indicated by this cross section, water-chemistry data and stable isotopes from veins on the south side of the fault show that little or no cross-fault CO2 migration is occurring. From HCO3 concentration in different wells (Heath, 2004), the CO2 gas may have infiltrated into many of the sandstone formations such as the Entrada, Navajo, Kayenta, and Wingate (schematic CO2-filled reservoirs are not shown to scale). Fractures related to the faulting allow infiltration of the CO2-charged groundwater through the otherwise sealing cap rock (arrows). Springs and geysers mark points where CO2-charged groundwater escapes along natural fractures, or through wellbores that penetrate the reservoirs.

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of the fault rocks at depth is partly dependent on the type of rocks that are juxtaposed across the faults and the amount of displacement that those fault rocks have undergone. The faults in this area offset a series of clean sandstones and shale-rich rocks and therefore could be expected to produce a clay-rich fault gouge that would be expected to act as a barrier to cross-fault flow. It must be emphasized that predicting fault seal from throw distributions is prone to error, and a small variation in fault-zone thickness or properties can create an apparent “hole” where fluid can leak through the fault. In contrast to the fault rocks, the shale-rich units that provide the topseal are leaking. Lithologically similar cap rocks have retained their integrity in CO2 reservoirs elsewhere in the Paradox Basin (e.g., McElmo Dome, Lisbon Dome); therefore, an explanation must be sought for why the cap rocks have failed at this location. Prior to drilling of the well, the leakage was focused in the immediate footwall to the faults. We suggest that fractures that formed in the cap rock as part of the damage zone to the faults are providing a conduit for leakage. It is also possible that an increase in CO2 volume at shallow depths leads to hydrofracturing, therefore enhancing fracture permeability. The fractures through the cap rock must have stayed open for substantial amounts of time (i.e., they are not self sealing). The strength and mechanical behavior of the cap rock units and the hydrodynamic behavior of CO2-rich fluid at shallow depths is poorly understood. Without such data, a reliable prediction could not presently be made of the integrity of cap rocks in similar structural settings. 3. Migration pathways. In our conceptual shallow reservoir model, the CO2-rich waters are sourced from the Wingate and Navajo Formations. Chemical analyses of groundwater from oil and gas exploratory and development wells, water wells and springs within ~100 km of the field area indicate that high dissolved CO2 concentrations are common in many formations from the Devonian Elbert Formation to the Jurassic Entrada Sandstone as well as the Navajo, Kayenta, and Wingate Sandstones (Heath, 2004). This distribution of CO2 content suggests that there is a sequence of stacked CO2-charged aquifers above the primary CO2 source. This is a critical issue for CO2 sequestration since leakage and migration of CO2 into formations overlying the intended storage formation may provide secondary trapping sites, reducing the overall flux to the surface.

At the outcrops along the Little Grand Wash fault, we looked at faults developed in mixed clastic sequences. The variety of structures that were developed in these faults was controlled by the lithology and lithological variation of the host rock. Today we will be looking at faults entirely within the aeolian Jurassic Navajo sandstones of the northern part of the San Rafael Swell. The primary variation of fault structure that we will see today is due to displacement variations on the fault. One way to infer the temporal evolution is by comparing faults that have different amounts of displacement, the idea being that faults with small displacements represent the early stages of development of a larger fault. Therefore, as we walk from low displacement sites to high displacement sites, we can imagine that we are looking at the growth of a fault within an aeolian sandstone. Road Log to Stop 2.1 Mileage Interval (Cumulative) Description 0.0

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Day 2: The Big Hole Fault and the Chimney Rock Fault Array Our second day examines faults in the Jurassic-Triassic Navajo Sandstone at the northern end of the San Rafael Swell. West-, northwest-, and southwest-striking normal faults cut the sandstone and have modest amounts (meters to tens of meters) of slip, and are analogs to faults that may partition hydrocarbon reservoirs. The field trip will consist of a walk along one such fault: the Big Hole fault, which has received a significant amount of attention including a combined field and drilling study (Shipton et al., 2002). If time permits, we will also examine the Chimney Rock fault array, which is a set of normal faults that reveal characteristics associated with fault interactions.

Depart Green River. Travel west on Main Street, and go west on I-70. Take Exit 156 onto U.S. 6 northbound. Turn left (west) onto the gravel road called the Green River cutoff. The road trends NW. Turn left (south) and go under the railroad. The road merges with the cutoff; take a hard right to the north. We are driving over the easily erodible Cretaceous Mancos Shale, which dips gently east here. This road is NOT recommended if there has been heavy rain. If heavy rain does occur while you are in the field area, there is an alternative route out to the west that should be taken. The road bends west and then southwest, through the Cretaceous Dakota Sandstone. Note the pebble conglomerate blocks of the Buckhorn Conglomerate here. Jurassic Morrison Formation rocks, dipping east. The road bends south. To our west, the rocks dip west. For the next mile, we travel along the axis of the Woodside anticline. This is a west-vergent fold, perhaps a fault bend-fold, above a small west-directed thrust. The road bends to the west, and we drive upsection through the Morrison and Dakota. We pass east-dipping Morrison Formation rocks, followed by the thin-bedded Summerville, Curtis, and Entrada Formations.

Anatomy of reservoir-scale normal faults in central Utah 4.8

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Intersection with dirt roads to the north and south. Turn south to get to the Big Hole fault. High-clearance vehicles are recommended from this point. Drive carefully along the unimproved road. The road ahead goes across the Big Hole wash on an old cement crossing. Park north of this crossing.

Stop 2.1: The Big Hole Fault This stop consists of an ~3 h walk west along the Big Hole wash. Bring water, food, and hat. Start off by walking west along the Big Hole wash to the first outcrop of the fault zone. Along the way, note the well-developed en echelon veins that are exposed in the Carmel Formation. Geologic Setting The Big Hole fault is the southernmost fault in the Chimney Rock fault array in the northern San Rafael Swell (Fig. 15). These dip-slip normal faults accommodated north-south extension. The San Rafael Swell is a broad east-vergent monoclinal flexure, which is the result of broad arching of Precambrian through Early Tertiary rocks above a basement fault, which was active during the Laramide Orogeny (50 Ma). Timing constraints on the age of faulting are poor, but extension parallel to the axis of the San Rafael Swell, perhaps in regions of maximum plunge, may be the cause of the faulting (Shipton, 1999). A crude overburden calculation, not accounting for compaction, suggests that the Navajo Sandstone experienced a lithostatic load of 40–80 MPa and temperatures of 45–90 °C during faulting. The Chimney Rock fault array and the Big Hole fault are excellent analogues for oil field and aquifer-scale structures (Fig. 16). Fault trace lengths range from 100 m to 4 km, displacements range from 0 to in excess of 30 m (Krantz, 1988; Cowie and Shipton, 1998). The

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faults intersect to create compartments of relatively undeformed host rock in the Navajo Sandstone (Fig. 16). The Jurassic Navajo Sandstone consists of planar to crossbedded very fine to fine-grained aeolian arenite of ~90%–95% quartz, 1%–6% feldspar, and 1%–3% clays. At this locality, the Navajo Sandstone is friable and weakly cemented by very low volumes of quartz overgrowths, iron oxide cements and/or calcite cement. Primary porosity ranges from 13% to 25% (Shipton and Cowie, 2001). Permeabilities to water for the undeformed Navajo Sandstone, measured at room temperatures and pressures, are typically in the 100–1000 mD range (Hood and Patterson, 1984) but can be over 4000 mD (Shipton et al., 2002). Currently, wastewater from coalbed methane operations is injected into the Navajo Sandstone northwest of the study area (Conway et al., 1997). The Navajo Sandstone is 137–151 m thick in the study area and is overlain by the fine-grained Carmel Formation limestone. The contact between the Carmel Formation and the Navajo Sandstone is the key horizon for measuring throw across the faults in the area (Krantz, 1988; Shipton and Cowie, 2001). The Navajo overlies the Jurassic Kayenta Formation. The Big Hole fault strikes N70°E, dips 64°N at the surface, and has pure dip-slip slickenlines. The fault and its surrounding deformation is exposed in the Jurassic Navajo Sandstone and offset of the overlying Carmel Formation is used to determine slip across the fault (Cowie and Shipton, 1998). The fault has a maximum of 29 m of slip and slip decreases approximately linearly toward each end, with the easternmost measurable slip of ~8 m recorded in the Big Hole Wash (Figs. 15 and 16). The 8 m Displacement Site This is the easternmost outcrop where the deformation around the Big Hole fault can be seen in the Navajo Sandstone. East of this locality the fault is poorly expressed at the surface. Linear extrapolation of slip on the fault (Shipton, 1999; Shipton and Cowie, 2001) from the point of 8 m of slip suggests the tip

Figure 15. Geologic map of the Bighole fault study area. On the left site we superpose the generalized geologic map is shown, with detailed maps of the Bighole and Blueberry faults shown on the right. Bh8, BH 17, BH 20, and BH 24 are stops along the Bighole fault, and the number indicates the amount of fault throw in meters.

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line would be as much as 800–1000 m from the location where 8 m of slip is measured. The air photo expression of a faultrelated linear structure in the Carmel Formation indicates that the fault probably extends to at least 500 m east of this point. The top of the Navajo Sandstone forms the base of the wash in the hanging wall, and is 8 m higher in the footwall. This value of throw is about the lowest that can be resolved on three-dimensional seismic data; note that none of the detail of the fault damage zone would be visible in three-dimensional seismic. The main fault zone slip-surface exposed in the base of the wash typically lies at the edge (or sometimes within) a pale

zone of dense deformation bands that has accommodated the majority of the strain across the fault. This is the fault core. Where it crosses the wash, the fault core forms a raised wall and consists of two to three main slip-surfaces bounding a complex structure of deformation bands that generally strike subparallel to the main fault. The fault core is surrounded by a damage zone that consists of clusters of deformation bands and occasional slip-surfaces. Note that in plan view this geometry could provide an effective barrier of baffle to horizontal fluid flow perpendicular to the fault, but may not provide much impediment to horizontal flow parallel to the fault.

Figure 16. Detailed geologic map of the Big Hole fault from Shipton and Cowie (2001). The map shows the distribution of the Carmel and Navajo Formations, and in a general way depicts the distribution of structures related to the Big Hole fault. (B) Inset stereonet snows the poles to the subsidiary faults in the hanging wall and footwall of the Bighole fault.

Figure 17. Outcrop map of the vertical wall south of the 8 m site. Thin lines are deformation bands, and thick lines are slip-surfaces. Grey areas are sand at the base of the outcrop. The Big Hole fault surface is off the left-hand side of this map, and the location along the fault is shown in the upper right (from Shipton and Cowie, 2001).

Anatomy of reservoir-scale normal faults in central Utah The footwall block is well exposed and gently sloping (between 40° and 80° toward the west) (Fig. 17). One main cluster of deformation bands is synthetic to the fault. One smaller synthetic cluster, and three antithetic clusters can also be seen. The main antithetic cluster contains many anastomosing but discontinuous slip-surfaces. The two other antithetic clusters to the south contain short segments of slip-surfaces. The short segments of slip-surfaces in the southernmost cluster are usually associated with local thickening of single deformation bands. This is a relationship that is often seen in deformation bands. The two sets of faults define lozenge-shaped blocks of relatively undeformed host sediment. Within the clusters, smaller lozenge-shaped blocks are preserved between individual deformation bands. Note that in the cross-sectional view this geometry could provide an effective barrier or baffle to horizontal fluid flow perpendicular to the fault, and to a lesser degree, vertical flow parallel to the fault. In the hanging wall of the fault you can see the two shut-in drill holes that penetrated the fault at depths of ~300 ft (Shipton et al., 2002). These holes were cored and wireline logs were acquired. 17 m Throw Site Continue walking west along the fault trace. Approximately 100 m west of the 8 m site, the fault is exposed on the north side of the wash, and ~300 m west along the fault strike from the last outcrop is a large flat outcrop exposed in the base of the wash. At this locality, the Big Hole fault consists of two strands. The total throw on the J2 unconformity in the cliff to the east is 17 m. The top of the Navajo Sandstone has been eroded from between the two strands, but the northern strand must have more than 3 m of throw and the southern strand must have less than 14 m (Fig. 18). Two outcrop-scale maps were made at Chipmunk Flat, one at each strand. The northern outcrop map is 9 m by 7 m (Fig. 18, left). A single slip-surface can be traced through the center of the fault zone. Other slip-surfaces anastomose around this main one defining a zone of concentrated deformation. The width of this zone is highly variable from just a single slip-surface up to 30 cm wide. In the western half of the map the deformation around the fault extends for ~1–1.5 m symmetrically around the fault. In the eastern half of the map the deformation around the fault is mostly within the footwall. To the north there are no clusters of deformation. The outcrop to the south is rather patchy due to sand trapped between the two upstanding fault strands. The southern outcrop map is 5 m by 18 m (Fig. 18, right). The southern strand has many anastomosing slip-surfaces. Intense arrays of deformation occur close to the fault mostly within the footwall. Complex clusters of single and multistrand deformation bands can be traced up to 14 m to the south. Clusters of deformation bands can be identified by groups of bands with local thickening, though it is hard to pick out individual clusters on the map. In both maps, the structures trend subparallel to the fault zone and define blocks of undeformed host rock between them.

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20 m Throw Site Continue west ~150 m along the fault strike from the last outcrop to another outcrop exposed in the base of the wash. The fault here has one main strand that forms a large step in the outcrop (Fig. 19). At this location the fault has 24 m of displacement. Almost all the deformation is concentrated on a single strand of the fault zone. Other discontinuous slip-surfaces anastomose around the main surface. This defines a zone of concentrated deformation about the main fault (fault core) of varying width along strike of between 10 and 30 cm. Within the main fault is a small pod of pale green glassy material ~1 m long and 3 cm wide. To the south of the main fault there is one large antithetic cluster that also forms a step in the outcrop. Good quality outcrop extends 180 m southwards and there are no more large clusters up to this point. To the north of the fault, several large clusters are seen up to 26 m from the fault. Only one of

Figure 18. Plan view outcrop maps of the fault strand at the 20 m site (from Shipton and Cowie, 2001).

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these clusters has a significant number of slip-surfaces in it. A wide variety of widths and numbers of bands can be seen in the clusters. Beyond the map the outcrop is mostly covered in sand, but no more deformation bands can be seen in the sandfree patches. At 150 m into the hanging wall of the fault, a large cluster can be seen, followed by several small clusters. Although the exposure is not good, the cliff face to the east shows that there are no deformation bands in the gap between the two clusters. This large cluster contains some of the highest deformation band density seen, but is not associated with any displacement of the top Navajo Sandstone. Summary of Bighole Fault Structure and Development The fault core of the Bighole fault is a zone of fine-grained, pale-colored rock up to 35 cm thick bounded by, or containing, one or more slip surfaces. The fault core generally has a resistant, glassy, almost mylonitic appearance though in thin section it can be seen that it contains only cataclastic deformation. Typically, several anastomosing slip surface strands are present in the main fault zone. Slip surfaces are highly polished or mineralized and

Figure 19. Cross section view outcrop map of the fault strand at the 24 m site. From Shipton et al. (2005).

contain slickenlines. Both in outcrop and in drill core, these can be tightly mated surfaces or act as planes of parting. In plan view, the through-going slip surface of the main fault plane is decorated with elongated “pods” of fault core (Fig. 20A). These lozenge-shaped pods are separated by areas with almost zero thickness of fault core. In cross section view, the fault core material is localized at intersections between synthetic and antithetic deformation band clusters that contain slip surfaces (Fig. 20B). Thus the fault core can be visualized in three-dimensions as being thickest at the intersection of synthetic and antithetic deformation-band clusters that are not totally planar, producing pods in plan view. In deformation band clusters away from the fault, the fault core can be developed (up to 10 cm thick) without slip surfaces (Figs. 20A and 20B). On average the offset across a zone of deformation bands is proportional to the number of bands in the zone (Aydin and Johnson, 1978; Mair et al., 2000). Narrow (<1 mm thick), planar, polished, slip-surfaces occur within and along the edges of clusters of deformation bands. Unlike deformation band zones, slip-surfaces can accumulate very large offsets (meters to hundreds of meters). Deformation bands have been shown to be effective hydrocarbon and water seals even at very small offsets (Antonellini et al., 1999). If slip-surfaces remain as open fractures at depth, they may present an enhanced permeability within a damage zone that otherwise may have very low permeability (Antonellini and Aydin, 1994; Shipton et al., 2002). Con-

Figure 20. (A) Model of Aydin and Johnson (1978) modified to show development of slip surface and fault core. (B) Slip surfaces nucleate at local points of high intensity grain crushing, (C) Slip surface segments propagate through developing zone of bands to link and eventually form through-going surface. (D) As displacement increases along a through-going slip surface, further development of is focused at kinematically incompatible linkage points. A thin veneer of highly comminuted fault core material lines each of the slip surfaces, but is too narrow to be represented on these figures. From Shipton et al. (2005).

Anatomy of reservoir-scale normal faults in central Utah sequently, understanding the evolution of slip-surfaces within networks of deformation bands is crucial for the prediction of fault zone permeability. Our data, acquired at a variety of scales, show that faults in aeolian sandstones consist of a four-component system (host rock, deformation bands, fault core and slip-surfaces), and that each component exhibits its own set of structural characteristics and hydraulic parameters. Even though we examined a “simple” system with one host rock and with one normal fault, significant variability in the fault zone permeability was observed. The host rock here may be one of four types, with permeabilities varying between 200 to >2000 md (Sections 3 and 5). The fault core has a relatively narrow range of permeability of typically less than 1 md (Section 5), but its thickness can vary dramatically over short distances sampled in the boreholes, and as seen in outcrop. Deformation bands likewise exhibit a relatively small range of variation in permeability, but their orientations and densities can also vary. Finally, the narrow slip-surface may be either a tight zone of low permeability or fractures may form along them, creating a zone of enhanced flow. Return east back to the vehicles. Depending on the time and interest of the group, an alternative stop after the Big Hole fault is the Chimney Rock fault array. Road Log to Stop 2.2 From the parking area, return north to the main Green River cutoff.

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Intersection with the main road; continue north. Fork in road—veer left. The road drops into a small wash. Park here or turn left onto the dirt track to a parking area.

Stop 2.2: The Blueberry Fault The Blueberry fault is 3.6 km long with a maximum throw of 30 m. The eastern tip of this fault is superbly exposed in a dry river canyon and detailed measurements of the throw profile at the fault tip are presented in Cowie and Shipton (1998). Figure 21 shows the map of the Blueberry fault tip exposed in the Blueberry fault Tip Canyon. As you follow the wash to the southwest it narrows down to a tight slot canyon, note the sparse deformation band clusters in the wall of this canyon. The canyon widens out again where it intersects the tip of the Blueberry fault. The southern side of the canyon is shallowly dipping, with patchy outcrop covered with some soil and vegetation. The subvertical northern wall is well exposed and it is possible to get a good view of the structures within it by scrambling up the southern wall. The Blueberry fault can be traced from the southern side of the canyon where it offsets the Navajo Sandstone. At this point the fault plane is a well-developed slip-surface with dip-slip and

Figure 21. Vertical outcrop map of the northern wall of the Blueberry fault tip canyon (displacement = 0 m). From Shipton and Cowie (2001).

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oblique-slip striations. The fault continues down the shallow southern slope and is buried in the sand at the base of the wash. Other deformation band clusters can be traced across the southern side of the canyon and are mostly synthetic or antithetic to the fault. Three clusters can be seen in the northern wall: two that dip to the south and one dipping to the north that is co-planar to the main Blueberry fault. At the top of the cliff, a single deformation band can be seen trending approximately co-planar to the fault. Slip-surfaces in the northern wall of the Tip Canyon are discontinuous. This confirms the observation that there is no throw of the J2 unconformity at the top of the north wall of the canyon. The damage zone is narrow around the Blueberry fault tip, but has a finite width at the point where the point of zero displacement is defined by the surveys of the J2 unconformity. The damage zone width around the Blueberry fault agrees with the scaling of damage zone width with throw seen around the Big Hole fault. Return to the vehicles. Return south to the Green River cutoff road, and either continue west, following the signs to, or return to the east along the road that we drove in on. If exiting to the west, drive west on the cutoff ~18 mi to an intersection with a graded gravel road. Take the gravel road north, which joins with State Route 10 near Huntington. Returning east, take the cutoff to U.S. Highway 6, and travel north to Price. REFERENCES CITED Antonellini, M.A., and Aydin, A., 1994, Effect of faulting on fluid flow in porous sandstones: petrophysical properties: American Association of Petroleum Geologists Bulletin, v. 78, p. 355–377. Antonellini, M.A., Aydin, A., and Orr, L., 1999, Outcrop aided characterisation of a faulted hydrocarbon reservoir: Arroyo Grande oil field, California, USA, in Haneberg, W.C., Mozley, P.S., Moore, C.J., and Goodwin, L.B., eds., Faults and subsurface fluid flow: American Geophysical Union Geophysical Monograph 113, p. 7–26. Aydin, A., and Johnson, A.M., 1978, Development of faults as zones of deformation bands and as slip surfaces in sandstones: Pure and Applied Geophysics, v. 116, p. 931–942, doi: 10.1007/BF00876547. Baer, J.L., and Rigby, J.K., 1978, Geology of the Crystal Geyser and the environmental implications of its effluent, Grand County, Utah: Utah Geology, v. 5, p. 125–130. Cappa, J.A., and Rice, D.D., 1995, Carbon dioxide in Mississippian rocks of the Paradox Basin and adjacent areas, Colorado, Utah, New Mexico, and Arizona: U.S. Geological Survey Bulletin 2000-H, 21 p. Chan, M.A., Parry, W.T., and Bowman, J.R., 2000, Diagenetic hematite and manganese oxides and fault-related fluid flow in Jurassic sandstones, southeastern Utah: American Association of Petroleum Geologists Bulletin, v. 84, p. 1281–1310. Conway, M.W., Barree, R.D., Hollingshead, J., Willis, C., and Farrens, M., 1997, Characterization and performance of injection wells in the Wingate and Navajo sandstones: Tuscaloosa, Proceedings, International Coalbed Methane Symposium, May 1997, p. 467–480. Cowie, P.A., and Shipton, Z.K., 1998, Fault tip displacement gradients and process zone dimensions: Journal of Structural Geology, v. 20, p. 983–997, doi: 10.1016/S0191-8141(98)00029-7. Demko, T.M., Currie, B.S., and Nicoll, K.A., 2004, Regional paleoclimatic and stratigraphic implications of paleosols and fluvial/overbank architecture in the Morrison Formation, (Upper Jurassic) Western Interior, USA, in Turner, C.E., Peterson, F., and Dunagan, S.P., eds., Reconstruction of the

extinct ecosystem of the Upper Jurassic Morrison Formation: Special Issue of Sedimentary Geology, v. 167, no. B, p. 115–136. Dockrill, D.L., Kirschner and Shipton, Z.K., 2006, Internal structure and evolution of fault-related active and ancient travertine deposits in East-Central Utah, U.S.A.: Sedimentary Geology (in press). Doelling, H.H., 1988, Geology of Salt Valley anticlines and Arches National Park, Grand County, Utah, in Doelling, H.H., Oviatt, C.G., and Huntoon, P.W., eds., Salt deformation in the Paradox region: Utah Geological and Mineral Survey Bulletin 122, p. 1–58. Doelling, H.H., 2001, Geologic map of the Moab and eastern part of the San Rafael Desert 30′ × 60′ quadrangles, Grand and Emery Counties, Utah, and Mesa, Colorado: Utah Geological Survey Map 180, scale 1:100,000. Doelling, H.H., 2002, Interim geologic map of the San Rafael Desert 30′ × 60′ quadrangle, Emery and Grand Counties, Utah: Utah Geologic Survey Open File Report 404, 22 p. Heath, J.E., 2004, Hydrogeochemical characterization of leaking carbon dioxide–charged fault zones in east-central Utah [M.S. Thesis]: Logan, Utah State University, 166 p. Hood, J.W., and Patterson, D.J., 1984, Bedrock aquifers in the northern San Rafael Swell area, Utah, with special emphasis on the Navajo Sandstone: State of Utah Department of Natural Resources Technical Publication no. 78, 128 p. Kennedy, B.M., Kharaka, Y.K., Evans, W.C., Ellwood, A., DePaolo, D.J., Thordsen, J., Ambats, G., and Mariner, R.H., 1997, Mantle fluids in the Sand Andreas fault system: California: Science, v. 278, p. 1278–1281, doi: 10.1126/science.278.5341.1278. Krantz, R.W., 1988, Multiple fault sets and three-dimensional strain: theory and application: Journal of Structural Geology, v. 10, p. 225–237, doi: 10.1016/0191-8141(88)90056-9. Mair, K., Main, I., and Elphick, S., 2000, Sequential growth of deformation bands in the laboratory: Journal of Structural Geology, v. 22, p. 25–42, doi: 10.1016/S0191-8141(99)00124-8. Pevear, D.R., Vrolijk, P.J., and Longstaffe, F.J., 1997, Timing of Moab Fault displacement and fluid movement integrated with burial history using radiogenic and stable isotopes, in Hendry, J.P., Carey, P., Parnell, J., Ruffell, A, and Worden, R., eds., Geofluids II: Chippenham, UK, Antony Rowe Ltd., p. 42–45. Peterson, F., 1984, Fluvial sedimentation on a quivering craton: influence of slight crustal movements on fluvial processes, Upper Jurassic Morrison Formation, western Colorado Plateau: Sedimentary Geology, v. 38, p. 21–49, doi: 10.1016/0037-0738(84)90073-3. Pipiringos, G.N., and O’Sullivan, R.B., 1978, Principal unconformities in Triassic and Jurassic rocks, Western Interior United States—a preliminary survey: U.S. Geological Survey Professional Paper 1035-A, 29 p. Shipton, Z.K., 1999, Fault displacement profiles and off-fault deformation: interpreting the records of fault growth at the Chimney Rock fault array, Utah, USA [Ph.D. thesis]: Edinburgh University, 251 p. Shipton, Z.K., and Cowie, P.A., 2001, Damage zone and slip-surface evolution over μm to km scales in high-porosity Navajo sandstone: Utah: Journal of Structural Geology, v. 23, p. 1825–1844, doi: 10.1016/S0191-8141(01)00035-9. Shipton, Z.K., Evans, J.P., Robeson, K.R., Forster, C.B., and Snelgrove, S.H., 2002, Structural heterogeneity and permeability in faulted eolian sandstone: Implications for subsurface modeling of faults: American Association of Petroleum Geologists Bulletin, v. 86, p. 863–884. Shipton, Z.K., Evans, J.P., and Thompson, L.B., 2005, The geometry and thickness of deformation-band fault core and its influence on sealing characteristics of deformation-band fault zones, in Sorkhabi, R., and Tsuji, Y., eds., Faults, fluid flow, and petroleum traps: American Association of Petroleum Geologists Memoir 85, p. 181–195. Turner, C.E., and Peterson, F., 2004, Reconstruction of the Upper Jurassic Morrison Formation extinct ecosystem—a synthesis, in Turner, C.E., Peterson, F., and Dunagan, S.P., eds., Reconstruction of the extinct ecosystem of the Upper Jurassic Morrison Formation: Special Issue of Sedimentary Geology, v. 167, p. 309–355. Williams, A., 2005, Structural analysis of the Little Grand Wash and Salt Wash faults to investigate the leakage of CO2 from a natural reservoir [M.S. thesis]: Logan, Utah State University, 92 p.

Printed in the USA

Geological Society of America Field Guide 6 2005

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths of the Henry Mountains, Southern Utah Sven Morgan Department of Geology, Central Michigan University, Mt. Pleasant, Michigan 48858, USA Eric Horsman Basil Tikoff Department of Geology and Geophysics, University of Wisconsin, Madison, Wisconsin 53706, USA Michel de Saint-Blanquat Guillaume Habert LMTG-UMR5563/Observatoire Midi-Pyrénées, CNRS/Université Paul-Sabatier, 14 av. Edouard-Belin, 31400 Toulouse, France

ABSTRACT Small intrusions (<3 km2) on the margins of the Henry Mountains intrusive complex of southern Utah are exceptionally well exposed in three dimensions and have a variety of shapes. Our examination of the geometry, structures, and fabric of the Maiden Creek sill, Trachyte Mesa laccolith, and the Black Mesa bysmalith (cylindrical intrusion bounded by vertical faults) suggests that this range of intrusion geometry may reflect a continuum of igneous emplacement as volume increases through magma sheeting. Intrusions begin as thin sills and through incremental injection of additional sheets, inflate into laccoliths. Marginal wall rocks are strained and rotated upward. Further sheet emplacement leads to the formation of a fault at the margin of the inflating intrusion. This fault accommodates piston-like uplift of the intrusion’s roof and results in the formation of a bysmalith. All three of these intrusions exhibit evidence for sheeting, although the evidence is weakest on the margins of the Black Mesa bysmalith. Solid-state shear zones exist between sheets in the Maiden Creek sill and on the margins of the Trachyte Mesa laccolith. Cataclastic zones also separate sheets within the Trachyte Mesa laccolith. Evidence for sheeting in the interior of the Trachyte Mesa laccolith is solely based on differences in weathering and jointing patterns. Evidence for sheeting on the margins of the Black Mesa bysmalith is based on the differences in lineation patterns and also on the distribution of cataclastic zones. Keywords: laccolith, Henry Mountains, sill, sheet, magma emplacement.

Morgan, S., Horsman, E., Tikoff, B., Saint-Blanquat, M.de, and Habert, G., 2005, Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths of the Henry Mountains, southern Utah, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 283–309, doi: 10.1130/2005.fld006(14). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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INTRODUCTION The intrusions of the Henry Mountains of south-central Utah provide an exceptional setting for the study of igneous emplacement processes. The igneous bodies intrude the welldocumented flat-lying stratigraphy of the Colorado Plateau and therefore displacement of the wall rocks resulting from emplacement of magma is well constrained (e.g., Gilbert, 1877; Hunt, 1953; Pollard and Johnson, 1973; Jackson and Pollard, 1988; Habert and de Saint Blanquat, 2004; Horsman et al., 2005). The intrusions are mid-Tertiary in age (Nelson et al., 1992) and therefore postdate the minor Laramide orogenic activity that affected this part of the Colorado Plateau. Consequently, fabric within the intrusions primarily reflects finite strain produced by magmatic flow during emplacement and lacks a significant concurrent or subsequent tectonic overprint. The Henry Mountains are the type locality to examine laccoliths since G.K. Gilbert originated the term laccolite (1877) based on his pioneering work here in the late 1800s. Gilbert’s mule- and horse-driven expeditions are themselves a significant accomplishment in physical endurance and scientific exploration, and well documented in Hunt’s reprinting (1988) of Gilbert’s field notes. Gilbert demonstrated that the igneous rocks of the Henry Mountains were actually intrusive (groundbreaking at the time) and he was one of the first to realize that magmas can deform their wall rocks. Gilbert envisioned a two-stage process of magma emplacement whereby the initial intrusion is sill-like, but with continued flow of magma, vertical growth is initiated and horizontal spreading ceases. The details of the growth of these intrusions have been debated ever since. It is the purpose of this trip to reexamine some of the same intrusions that Gilbert studied and to illustrate that at least in some intrusions, abundant evidence exists for the growth of the intrusion by the stacking of multiple magma sheets. Hunt (1953), after years of detailed mapping in the Henry Mountains in the 1930s (also on horseback), reinterpreted the five main intrusive centers and/or mountains as stocks, not laccoliths. The principal difference being that a laccolith has a flat lying sedimentary “floor,” and a stock continues downward with depth. Hunt (1953) agreed with Gilbert in that the smaller surrounding intrusions were laccoliths and were fed from the large intrusive centers. Detailed mapping and geophysical work by Jackson and Pollard (1988) led to an interpretation more in agreement with Gilbert (1877); i.e., the five main intrusive centers were “floored” laccoliths and not stocks. This field guide concentrates on the small (<3 km diameter) intrusive bodies that occur on the east side of the Mount Hillers intrusive center: the Maiden Creek sill, Trachyte Mesa laccolith, and the Black Mesa bysmalith. We will also briefly examine the Sawtooth Ridge intrusion. We call these small intrusions satellites, because they are located between 5 and 10 km from the closest intrusive center (Mount Hillers), their magma volumes are small relative to the main intrusive centers, and they are radially distributed about the intrusive centers. The elongate shape of

Trachyte Mesa is also oriented along a line that intersects Mount Hillers, and magmatic lineations within the Trachyte Mesa laccolith are also subparallel to this line. These satellite intrusions, on a much smaller scale, resemble the intrusive centers in composition, shape, and style of emplacement (contact geometries are similar). In contrast to their larger neighbors, the satellites are much better exposed and more readily accessible (low elevation and proximity to roads). All are close to Utah Highway 276 and to dirt roads, which allows visitation of all these intrusions in one or two days. Our work is different than previous studies in that we have focused on the fabrics within the intrusions, as well as the structures and geometry of the intrusions. In particular, we have focused on the role of discrete magma pulses, observed in the field as magma sheets, in constructing these igneous bodies. We hypothesize that these three intrusions—the Maiden Creek sill, Trachyte Mesa laccolith, and the Black Mesa bysmalith (from smallest to largest)—may reflect the evolution of a magma chamber with increasing magma input. Alternatively stated, the Black Mesa bysmalith may have originated as a sill (a Maiden Creek “phase”) that became a laccolith (a Trachyte Mesa “phase”) before evolving into its present bysmalith form (a bysmalith is a cylindrical intrusion with vertical faults as contacts, whereas a laccolith is more dome-like). We can show many rock exposures that illustrate the multiple sheet-like construction of the Maiden Creek Sill and Trachyte Mesa laccolith. We also have evidence, although weaker, for slightly more cryptic sheets within the upper portion of the Black Mesa bysmalith. GEOLOGY AND HISTORY OF THE HENRY MOUNTAINS The Henry Mountains of south-central Utah are a 90-kmlong and 30-km-wide Tertiary igneous complex on the Colorado Plateau (Fig. 1). Although they are the largest of seven laccolithic ranges found on the Colorado Plateau (Stokes, 1988), they were one of the last surveyed and the last-named ranges in the lower 48 states. John Wesley Powell called them the Unknown Mountains on his first trip down the Colorado River in 1869 (Kelsey, 1990). Powell named them in 1871–1872 after Joseph Henry, a close friend and secretary of the Smithsonian Institution, on his return voyage down the Colorado River. The first non-native set foot in the Henrys in 1872, when A.H. Thompson, a geographer from Powell’s second party, explored the northern peaks; Mount Ellen is named after Thompson’s wife (Fillmore, 2000). In 1875 Powell assigned Grove Karl Gilbert to study the “volcanic” mountain range, and Gilbert made two trips (two weeks in 1875 and two months in 1876) (Fillmore, 2000). The range is composed of five distinct igneous centers that were emplaced into basically flat-lying stratigraphy (regional dips of 1–2° W). Ten kilometers to the west, the stratigraphy is abruptly upturned in the Waterpocket Monocline, and the canyons of the Colorado River are <10 km to the southwest. The peaks attain heights of over 11,000 ft, while the elevation of the

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths surrounding flat-lying rocks of the plateau hovers around 4000– 5000 ft. A striking aspect to the Henry Mountains, besides their size and nature of origin, is the abruptness with which the mostly Mesozoic wall and roof rocks rotate upward on the flanks of the intrusions to produce these mountains; this is readily observed from Hwy 95 and Hwy 276, which parallel the mountain chain. Even though the mapped intrusives have steep margins, it is evident based on the much more subtle slopes of the immediately surrounding sedimentary layers that these intrusions expand over a much greater area than what is actually exposed. Mount Hillers is the best exposed of the five Henry Mountains intrusive centers and presently forms a dome with a diameter of ~15 km. Several steep, narrow canyons cut into Mount Hillers and expose cross sections through numerous intrusions radiating away from the center of the dome. The sedimentary rocks on the top have also been eroded away partially exposing the igneous core. Assuming that there were sedimentary rocks on top of Mount Hillers, the vertical displacement of the sedimentary rocks is ~2.5 km (Hunt, 1953; Jackson and Pollard, 1988). The sedimentary section is dominated by sandstones and shales that range in age from Permian to Cretaceous and is ~2.7 km thick (Peterson et al., 1980; Jackson and Pollard, 1988). Jackson and Pollard (1988) determined that at the time of intrusion, the intrusives in the Henry Mountains were buried by 3–4 km of sedimentary overburden. FIELD TRIP Overview of Day 1

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Figure 1. Regional map showing the area of study (small rectangle), Stop 1.1, and the five intrusive centers and peaks of the Henry Mountains, Utah. Modified after Habert and de Saint Blanquat (2004).

Directions to Stop 1.1 Go east (and south) on Utah Hwy 95. Mile marker 2 (while driving by). At two o’clock (treating the car, in map view, as a clock, the front of the car is twelve o’clock, passenger side is three o’clock, etc.) is the Mount Ellen intrusive center. This is one of the five main intrusive centers in the Henry Mountains. There are two prominent mountains located at the periphery of Mount Ellen. Table Mountain is a mesa on the north end of the intrusive complex, and Bull Mountain is a more cylindrical looking mountain on the east side of the intrusive complex. Both mountains are bysmalith intrusions similar to Black Mesa, which we will observe later. Mile marker 13 (while driving by). Shallowly dipping striped sediments E of here are the Carmel, Summerville, and Morrison Formations. The desert floor we are driving on is the Entrada Sandstone. Mile markers 17–18 (while driving by). Reddish, eolian sediments of the Entrada Formation in a small canyon.

Mile marker 26. Reset odometer to zero. Turn right onto Utah Hwy 276 toward Ticaboo. Mile markers start at 0 at the intersection with Utah Hwy 95. Drive 1.7 mi to a small rise in the road and pull over on the side of the road for Stop 1.1. Stop 1.1: Overview of Henry Mountain Intrusive Centers The five intrusive centers of the Henry Mountains are visible from this point (Fig. 1). Looking due south along the road (and calling this 12 o’clock), Mount Holmes is at 12, Mount Ellsworth at 12:30, Mount Hillers at 1, Mount Pennell at 3, and Mount Ellen at 5. Each of these intrusive centers is composed of many component igneous intrusions. Mounts Ellsworth and Holmes are the smallest intrusive centers, and the exposed igneous rocks are composed of dikes and sills and a few relatively small laccoliths. Sedimentary rocks dominate the margins and can be traced almost to the top. The sedimentary layering is steeply inclined, indicating that more igneous rocks are beneath. Mount Ellen is the largest of the three northerly mountains and all three

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are composed of a great many intrusions with a wide range of sizes and geometries. The sedimentary layering is best exposed at the base of these intrusions and intermittently exposed through the middle elevations at various dips. On the east flank of Mount Hillers, the Black Mesa and Sawtooth Ridge intrusions are visible on the skyline (Fig. 2). A sheer cliff forms the east side of the Black Mesa bysmalith and Sawtooth Ridge is noted by its prominent jagged edge against the skyline. Directly in front of us (middle ground), the Trachyte Mesa intrusion is exposed. The Maiden Creek sill is exposed just east of Black Mesa. Our strategy is to visit the least-developed intrusions first (i.e., the earliest stages in a possible continuum of development of an inflating magma chamber and/or intrusion), and visit the more-developed intrusions later. From here we visit the Maiden Creek sill for Stops 1.2–1.8, break for lunch, and then spend the afternoon at the Trachyte Mesa laccolith for Stops 1.9–1.13. Directions to Stop 1.2 Continue south on Rt. 276. At mile 6.7, turn right onto dirt road. This is immediately (<100 m) before mile marker 7 on

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Utah Hwy 276. At mile 9.7, park on side of dirt road where road turns SE and there is some room to park in the bend. Walk ESE down the stream valley for ~0.5 km (Fig. 3) until igneous rocks are seen in the stream bottom. Part I: The Maiden Creek Sill Stop 1.2: Composition and Top Contact of the Maiden Creek Sill GPS (UTM): 536090, 4195836, zone 12, NAD27. Main points: (1) Composition and texture of the Maiden Creek sill. (2) The geometry of the body is a thick (~25 m) sill. (3) The top contact contains solid-state fabric in the uppermost ~3 cm, the remainder of the intrusion contains magmatic fabrics. Lineation and foliation are generally parallel in both fabric types. This stop provides a good spot to view the basic cross-sectional geometry of the Maiden Creek sill (Figs. 4 and 5). The planar-horizontal top contact of the intrusion with the overlying Entrada Sandstone is well exposed here. We will soon examine the planar-horizontal bottom contact of the intrusion in the steepwalled gorge ~150 m east of here. The intrusion is generally ~25 m thick but thins toward its margins in some areas. The horizontal, concordant top and bottom contacts of the intrusion lead us to call this a sill. However, the map view geometry of the intrusion is considerably more complex, as will become apparent shortly. This stop also provides an introduction to the igneous rock of the Henry Mountains, which is a mostly homogeneous plagioclase-hornblende porphyry in which phenocrysts of feldspar and amphibole lie in a very fine-grained gray matrix (Fig. 6; see also Hunt, 1953; Engel, 1959; Nelson et al., 1992). In the Maiden Creek sill, feldspar phenocrysts make up 30%–35% of the rock by volume and are generally euhedral laths 0.2–1 cm in diameter. Amphibole phenocrysts make up 5%–15% of the rock by volume and are euhedral needles 0.1–0.5 cm in length. Other phenocrysts include euhedral to subhedral oxide grains, which generally have a maximum diameter of 0.2 mm and make up <2% of the porphyry by volume, and apatite and sphene, both of which are generally euhedral and <1% by volume. The very fine-grained matrix generally makes up 50% or more of the rock and is composed of microcrystalline feldspar, amphibole, and oxides. Because of the well-exposed upper contact of the intrusion at this location, we can examine the relationship between the contact and fabric development in the igneous rock. The outermost ~3 cm of the intrusion generally have well developed solidstate fabric, which is defined here by the cataclastic stretching of feldspar and amphibole phenocrysts (Fig. 6C). These deformed phenocrysts define a prominent lineation. More than ~3 cm from the contact, solid-state fabric is essentially nonexistent and magmatic fabric is developed instead. This magmatic fabric is defined by the preferential alignment of undeformed phenocrysts, particularly elongate amphibole crystals (Fig. 6B). In locations where both solid-state and magmatic fabrics are exposed, foliation and lineation orientations in both fabric types of fabric are generally

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parallel. These same observations will be seen on the margins of the Trachyte Mesa laccolith. Directions to Stop 1.3 Head NE down the N side of the creek (Fig. 3) for ~100 m until you are on relatively flat-lying sedimentary rocks. You should have an overview (looking S across the streambed) of the N-facing side of the stream canyon and ~4 m high towers on the N side of the streambed. Stop 1.3: Lateral Terminations GPS (UTM): 536186, 4195897. Main points: (1) There are two bulbous terminations on the lateral margin of the sill, suggesting different magma sheets. (2) There is upward tilting of wall rocks on the margin of the intrusion. (3) This margin belongs to an NS-elongated part (finger) of the Maiden Creek sill. Looking S from this vantage point, the lateral contact of the intrusion with the surrounding Entrada sandstone is beautifully exposed (Fig. 7A). The contact is complex; it is composed of two vertically stacked bulbous igneous terminations. As we will observe at other locations on the intrusion, this geometry is typical of most exposed lateral contacts of the Maiden Creek sill. We suggest that this geometry is most easily produced by the

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stacking of separate igneous sheets, and two separate sheets can be observed at Stop 1.8. The relationships between the sedimentary host rock and the intrusion are better preserved at this location than at most others on the Maiden Creek sill. The sediments here tilt upward above each of the upward-facing portions of each bulbous sheet termination (Fig. 7A). This observation indicates that at least some of the host-rock displacement necessary for emplacement of the intrusion is accomplished through tilting of the sediments. The lateral margin exposed here forms the eastern edge of a NS-elongated portion of the Maiden Creek sill (which we will refer to as a finger, following Pollard et al., 1975) that projects out from the main body of the intrusion. The ridge immediately E of here is defined by another NS-elongated finger. The lateral contact we have been considering and others in this area require that the current geometry of the fingers is very similar to the original emplacement geometry, i.e., these fingers were not formed by erosion of a once-contiguous sheet. Directions to Stop 1.4 Head (~100 m) ESE down to the streambed and continue through Secret Nap Spot gorge until you encounter a 10 m drop along the streambed.

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Stop 1.4: Cross Section in Secret Nap Spot Gorge GPS (UTM): no GPS measurement is possible here. Main points: (1) Complete cross section of one finger. (2) Bottom contact is conformable and not deformed. (3) No stoping visible in the igneous rock. Exposed at this stop is a complete cross section through one of the fingers of the Maiden Creek sill (E side of Fig. 5C). Although overlying sedimentary rocks are not clearly exposed atop the cliff faces here, they are in place nearby, as we will observe en route to a later stop. Examining the cliff faces, it is clear that there are no xenoliths of sedimentary rock within the intrusion. Sedimentary xenoliths are exceedingly rare in all of the intrusions we will be examining. This suggests that stoping is not an important space-making mechanism in these intrusions. Also exposed here is the concordant horizontal bottom contact of the Maiden Creek sill with the underlying Entrada Sandstone (here a shale). The bedding in the host rock below the contact is essentially undeformed. Clear evidence for three fingers (like this one) exists on the margins of the Maiden Creek sill and strong evidence for a fourth

finger (which we were atop at Stop 2) exists. Preserved contacts demonstrate that the finger-like lobes project out from the main body of the intrusion and are not merely erosional remnants of a main body that was once larger and has been dissected by streams. No clear textural boundary exists between the main body of the intrusion and the finger-like lobes. Each finger is 200–400 m long and is distinctly elongate with respect to both its cross-sectional thickness and map view width (Figs. 5A and 5B). In longitudinal cross section (Figs. 5C and 5D), each finger thins progressively away from the main body. The finger shown in Figure 5A thins from ~30 m thick near the main body to ~7 m thick over a distance of ~400 m. This thinning occurs as the base of the intrusion cuts up through the sedimentary section while the top of the intrusion resides at a consistent bedding-parallel stratigraphic level. Directions to Stop 1.5 Backtrack out of the gorge and then walk ~175 m NW up a smaller canyon developed in sedimentary rocks. Walk up the creek bed until igneous rock is encountered.

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Stop 1.5: Igneous-Sedimentary Contact on a Lateral Termination GPS (UTM): 536186, 4196016. Main points: (1) Solidstate deformation occurs locally on lateral contact. (2) Very little deformation or metamorphism occurs in wall rocks. The lateral contact of the intrusion with the host rock is exposed at this stop. Solid-state fabric in the outermost few centimeters of the intrusion is prominently developed. In general, the solid-state fabric patterns are complex and probably reflect the nature of the emplacement process, which accommodated both lateral and vertical expansion and longitudinal propagation of the sheet. As we observed at Stop 1.2 atop the intrusion, fabric patterns at the top (and bottom) contacts tend to be more consistent and predictable than those observed at the lateral contacts. This difference is presumably related to the more complex flow present in the bulbous lateral sheet terminations than that present at the planar top and bottom margins of the sheet. Similar relationships are seen at Trachyte Mesa laccolith, i.e., intense deformation at the lateral margins and little to no deformation at the bottom and top margins.

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Directions to Stop 1.6 Walk several hundred meters NE (Fig. 3) until cliff is encountered (edge of the intrusion). Notice along the walk the in situ horizontal sediments atop the intrusion. Stop 1.6: Geometry of Intrusion GPS (UTM): 536426, 4196131, Main points: (1) The Maiden Creek intrusion consists of 100-m-scale, finger-like intrusions that emanate from a central intrusion. (2) The top of the intrusion is relatively flat and concordant. The bottom of the intrusion cuts up-section, resulting in thinning of the intrusion along its length (longitudinal axis). (3) Bulbous terminations are stacked upon one another, on lateral margins of main part of the sill, suggesting that different magma sheets exist throughout the area of the intrusion. (4) Meter-scale topography exists locally on the top of the intrusion. From this vantage point, much of the eastern side of the Maiden Creek sill is visible. We will discuss several noteworthy features of the intrusion. First, looking S from here, the southeastern finger of the intrusion (which we looked at in cross

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C Figure 6. Microstructures in the Maiden Creek sill. (A) Field photo of solid-state fabric developed at contact. (B) Magmatic fabric. Photo is ~1.5 cm across. (C) Solid-state fabric. Photo is ~0.5 cm across. White streaks represent plagioclase phenocrysts that have been cataclastically altered to a much finer grain size and extended along the foliation. After Horsman et al. (2005).

section at Stop 4) can be clearly seen. The top of the finger is relatively flat and concordant, and flat-lying sediments are still in place in several locations atop this part of the intrusion. Although not clearly visible from here, the bottom contact of the finger cuts up-section along its length, resulting in thinning of the intrusion. Each of the fingers of the Maiden Creek sill has a similar geometry. These fingers radiate from the main body of the Maiden Creek sill, resulting in a complex map view geometry (Fig. 4). The main body of the intrusion (which we are now atop) is roughly elliptical in map view and has a relatively simple silllike geometry in cross section. This region of the intrusion is consistently 30–40 m thick. At least four separate, finger-like lobes project out from this main body. Each of these fingers is ~30–40 m thick where it begins to project out from the main body and thins progressively to a thickness of 5–10 m as distance from the main body increases (Fig. 5). These fingers extend 200–400 m out from the main body of the intrusion and are distinctly elongate with respect to both their map view width and their cross-sectional thickness. The numerous extant lateral contacts with sedimentary host rock strongly suggest that the current map pattern (Fig. 4) of the intrusion is very similar to the original intrusive geometry. This complex geometry may be a common characteristic of sills in general (e.g., Marsh, 2004) because the lateral margins of most sheet intrusions are rarely as well preserved as those seen on the Maiden Creek sill. Also visible from this vantage is the steep eastern side of the sill. Along this cliff the intrusion is 30–40 m thick and has two well-exposed vertically stacked bulbous lateral terminations (Fig. 7B). The lateral contacts with adjacent sedimentary rocks are locally preserved. As discussed earlier, these bulbous lateral terminations suggest that at least two magma sheets are stacked upon one another throughout the intrusion. These two sheets can be observed individually, as they thin to the NE and separate into individual sheets for several tens of meters (NE end of Fig. 5A; Stop 1.8). These intercalated sedimentary layers are observed in a few places and are dominantly found halfway between the base and the top of the sill, at the same location where the two bulbous terminations meet. Rare igneous-igneous contacts can also be traced from these screens of wall rock (Fig. 7C). These internal contacts are marked by a 1–2 cm zone of intense foliation and solid-state deformation. We will not visit the outcrop illustrated in Figure 7C, because one must climb down this cliff face, but we will see the two separate sheets from the road on Stop 1.8. One final feature visible from here is the locally developed m-scale topography of the top surface of the sill. Nascent dikes extending up from the roof of the intrusion occur in two places (one visible just W of here, the other on the NW exposed corner of the intrusion). Also present are m-scale ridges, generally oriented parallel to the strike of the edge of the intrusion. Johnson and Pollard (1973) interpreted similar ridges atop the Trachyte Mesa intrusion to record the position of the edge of that intrusion at various stages of progressive emplacement.

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Directions to Stop 1.7 Walk 200–300 m N, essentially on the top of the intrusion. Head for the overlook of a gully between two prominent NESW–oriented ridges (Fig. 3). Stop 1.7: Emplacement of the Maiden Creek Sill GPS (UTM): 536494, 4196351. Main points: (1) Magmatic flow is subparallel to the fingers. (2) Areal extent of the first sheet seems to have controlled the lateral extent of the second sheet. (3) Emplacement occurred primarily by upward vertical movement of the roof rocks although there are still space problems. The gully below us to the N is bordered by two ridges, each of which is held up by a finger of the Maiden Creek sill. Excellent lateral contacts of both fingers are preserved at the SW end of this gully (see Figure 15 in Pollard et al. [1975] for a photograph of a portion of one of these lateral contacts). Because we believe the current outcrop pattern corresponds to the original shape of the sill, an ideal opportunity exists to study the relationships between igneous fabric, emplacement

processes, and intrusion geometry. Fabric results described here are summarized from Horsman et al. (2005), who used several techniques to analyze the fabric within the sill. Throughout the intrusion, solid-state fabric is confined to the outermost ~3 cm of the intrusion (and the few internal contacts adjacent to the screens of wall rocks). At distances greater than ~3 cm from the contact with the sedimentary host rock, magmatic fabric is almost exclusively developed. The boundary between the regions of solid-state and magmatic fabric is gradational over 1–2 cm. Magmatic fabric is a more reliable tool than solid-state fabric for studying magmatic flow during emplacement because the solid-state fabric records both magmatic flow and the interaction of the magma with the wall rock. Solid-state fabric is consequently more complex and difficult to interpret than magmatic fabric, which reflects primarily finite strain produced by the late-stage magmatic flow during emplacement (Horsman et al., 2005). With these caveats in mind, the fabric results for the Maiden Creek can be considered.

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Magmatic lineation in a finger of the sill is generally subparallel to the elongation direction of the finger (Fig. 8). Field-measured lineations and AMS (anisotropy of magnetic susceptibility) lineations are generally subparallel (for the details of the fabric and a discussion of the AMS see Horsman et al., 2005). Magmatic foliation in the fingers is generally sub-horizontal. Consistent magmatic fabric patterns are also observed in the main body of the sill, where magmatic lineation tends to be oriented roughly radial and foliation is generally subhorizontal. We interpret these patterns to record general flow of magma away from an unexposed source region to the W into the main body and then out of the main body and into each finger. Although, we cannot distinguish whether the fingers are late ancillary intrusions fed by the main body, or if the main body is a region that has coalesced into a sheet of magma by spreading between early fingers (e.g., Pollard et al., 1975). We infer that the Maiden Creek sill consists of two separate, sequentially emplaced sheets that were emplaced as different pulses of magma. The two sheets have almost the exact same geometry and extent. Their similar cross-sectional geometry allows us to conclude that the pulses of magma were essentially identical

in volume. This observation suggests that the emplacement and extent of the first sheet controlled the extent of the second sheet. We envision the following scenario. The emplacement of the first sheet produces a weak, hot region surrounding the igneous body. During intrusion of subsequent igneous sheets, the adjacent heated sedimentary rock will be relatively weak and will fail before nearby cooler sediments. Consequently, after the emplacement of the first sheet, subsequent pulses of magma from the same feeder system intrude immediately adjacent to the previously intruded sheet. By this process, a thick sheet of igneous rock with a complex lobate three-dimensional geometry is built. The consequence is that the amount of magma intruding at any one time remains small. Uplift of overlying roof rocks is the dominant space-making mechanism. However, when viewed in cross section, there does not seem to be enough deflection of the layers on the margins (layers that rotate up at the contact with the Maiden Creek sill) to accommodate the amount of vertical space required by the thickness of the sill. This observation indicates that other mechanisms, such as lateral displacement and strain of adjacent host rocks (as

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Figure 8. Macroscopic lineations and magnetic lineations (insert) on top of the Maiden Creek sill. AMS—anisotropy of magnetic susceptibility. Modified after Horsman et al. (2005).

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horizontal sedimentary block separates the two fingers (Fig. 5A). To the W of this intercalated sedimentary block, within the one thick (stacked) finger, a 2–3 cm thick, solid-state, igneous-igneous contact can be traced from the edge of the block for several tens of meters and define the boundary between the lower and upper fingers. Immediately to the E of the termination of these two fingers, another very thin sheet appears, at a topographic level intermediate between the upper and lower fingers and extends for several tens of meters before thickening at the margin of the mesa. We suggest that the separation of one thick finger into an upper and a lower finger, as well as the internal contact to the W of this split, supports our hypothesis that the Maiden Creek sill was constructed by stacking two sheets that share the same areal extent.

we will observe at the Trachyte Mesa intrusion) may be responsible for some of the additional space making. Directions to Stop 1.8 Head W (contouring) until the road is seen. Walk to road and then southward to car. Turn cars around and head back N and E on dirt road until intersection with Utah Hwy 276. Set odometer to zero and turn right (S) on Hwy 276. Drive ~0.6 mi and park on the side of the road (Fig. 3). Look to the W toward the top of the mesa that terminates close to the road. Stop 1.8: Separate Fingers of the Maiden Creek Sill Main point: (1) Separation of the Maiden Creek sill into an upper and lower finger supports our model of two pulses of magma sheeting, one stacked upon the other. Looking up at the top of the mesa, from the W to the E, the NE finger of the Maiden Creek sill splits into an upper and a lower finger. These two thinner fingers extend for several tens of meters before terminating at the same lateral extent, and they both terminate along similarly oriented faults. A several-meter-thick,

Part II: Trachyte Mesa Laccolith Directions to Stop 1.9 Set odometer to zero. Turn around and head back N on Utah Hwy 276. At mile 1.6, just beyond (N) the rise in the road, turn left and drive SW onto a dirt road (Fig. 9). At mile 2.0, stop

Quaternary (?) Qa dominantly alluvium Tertiary Tp porphyry (Oligocene) Jurassic Js Summerville Fm Je dominantly Entrada Fm (some alluvium)

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Figure 9. Topographic map of Trachyte Mesa laccolith with stops indicated.

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on the dirt road when it approaches within ~40 m of a gully north of the road (Stop 1.9). Looking ~NW across the gully, there is a clear exposure of a meter-thick sandstone bed intercalated with igneous sheets at the top margin of the Trachyte Mesa laccolith. Stop 1.9: View of the Southeast Margin of the Trachyte Mesa Laccolith GPS (UTM): 536703, 4199129. Main point: Intercalated beds of Entrada Sandstone and igneous sheets within the margin of the Trachyte Mesa laccolith. This is an excellent view of the sheeted nature of the Trachyte Mesa laccolith and the subhorizontal basal contact (Fig. 10). The cliff face below the intrusion is the red sandstone of the Entrada Formation. The thickest and lowest sheet of the Trachyte Mesa laccolith is 8 m thick here. Directly above this basal sheet is a 1m-thick bed of Entrada sandstone. Directly above this sandstone layer are at least two more sheets. These two upper sheets have distinctively different erosional morphologies. Directions to Stop 1.10 Continue SW on dirt road for ~0.2 mi. Park at clearing by large juniper tree. GPS (UTM): 536254, 4198926. The road becomes less traveled beyond the juniper tree. At this point, the field trip will continue by foot. Head in direction 330 (azimuth) directly toward Bull Mountain (the last peak along the range moving toward the NE) on the E flank of Mount Ellen. Continue for ~300 m until you intersect streambed. Walk downstream until the bed of the stream is igneous rock.

south side, Mt. Ellen

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Figure 10. Photo of sandstone bed intercalated with igneous sheets, SE margin of Trachyte Mesa laccolith. Basal sheet is 8 m thick.

Stop 1.10: Top Contact of the Trachyte Mesa Laccolith GPS (UTM): 536086, 4199194. Main points: (1) The composition and texture of the Trachyte Mesa laccolith is very similar to the Maiden Creek sill. (2) The upper igneous contact with undeformed and unmetamorphosed sandstone is well exposed. (3) The sandstone layer on the top contact here can be traced over much of the top of the laccolith. This outcrop is assumed to be very close to the southern extent of the laccolith. The top contact is exposed and dips to the SW beneath sedimentary strata of the Entrada Formation. The thickness of the intrusion increases greatly from SW to NE along the streambed as the top contact rises 20 m over a lateral distance of 100 m. The base of the intrusion is not exposed, but it is assumed to be very shallow (<2–3 m) beneath the southwesternmost exposure of igneous rock. We assume this because the basal contact is exposed to the NE, outside of the stream gorge, along the cliff that was viewed from the road driving in. The top several cm of the contact of the Trachyte Mesa laccolith has been eroded at this location. At other locations where the actual top of the intrusion has not been eroded, there is a 2– 3 cm thick carapace of solid-state deformation. Very limited (several cm or less) erosion occurred here, evidenced by the upper contact with the sedimentary rocks exposed on the hillside. The texture and composition of the Trachyte Mesa laccolith is very similar to the Maiden Creek sill. The rock has a microgranular porphyritic texture with euhedral phenocrysts (up to 7 mm) of hornblende and plagioclase. The groundmass consists dominantly of microlaths of plagioclase and oxides (mostly magnetite). Most of the plagioclase phenocrysts exhibit concentric zonation. Many of the hornblende phenocrysts and larger oxides are partially or completely altered to finer-grained oxides and calcite. A few pyroxene and apatite phenocrysts are observed. The sandstone at the upper contact is unmetamorphosed and undeformed. Although, within several cm from the intrusion, we often observe pea-sized nodules within the sandstone that resist weathering. We can observe this feature here at the base of the sandstone layer. These nodules of sandstone have their pore spaces filled with calcite cement. Many of the oxides and hornblende in the igneous rock have also been replaced with calcite, documenting the fluids that passed through these rocks syn- to post-emplacement. The sandstone layer is full of cm-scale holes and larger cavities. This spheroidal weathering seems to be a marker for this particular sandstone layer. We can trace this weathered ~1-mthick sandstone layer over the top of the Trachyte Mesa laccolith and over to the NW margin. Directions to Stop 1.11 Walk downstream for 150 m, through a gorge that turns to the SE and just past a 1 m drop in the creek bed and continue onto two large (~4 m) boulders for an overview. Stop 1.11: Differential Erosion of Sheets GPS (UTM): 536236, 4199158. Main points: (1) There is geomorphic evidence for multiple intrusive sheets. (2) The

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths sedimentary layer beneath the intrusion is a massive sandstone layer. (3) The basal contact is concordant and climbs upward to the SE because the sedimentary beds are tilted down to the NW. The walls of the gorge we just walked through weather differently at different elevations. The upper part of the gorge, on both sides, is eroding and producing vertical cliff faces with multiple vertical joint faces. Approximately halfway down, both walls of the gorge step toward the center of the gorge. Several meters downward, the gorge walls become vertical once again. This differential erosion is best expressed on the N side of the gorge, where the jointing is also different between the top, middle, and bottom sections of the gorge walls. The middle section, which resists erosion and where there are very few joints, also increases in thickness gradually from the W to the E (down stream) and resembles a sill. We interpret this erosional profile as resulting from different igneous sheets. The differences in jointing within each “step” down the walls of the gorge are at least suggestive that these “steps” cooled at different rates or at different times. Basalt flows are routinely differentiated using differences in erosional profiles and differences in jointing. Cañón-Tapia and Coe (2002) were able to differentiate several different basalt flows along the Columbia River using AMS, which supported their identification of the flows based on textural, jointing, and erosional differences. However, we have found no textural evidence along the walls of this gorge for different sheets. Our AMS data (unpublished and not shown here), taken along multiple vertical profiles along the walls of this gorge, does not reveal any differences that may indicate distinct sheets. At the E edge of the gorge, the stream cuts through the margin of the laccolith. Where the streambed changes from igneous to sandstone is a small (<1 m) cliff. Standing on the edge of the cliff, the basal contact is not exposed on the cliff wall below, but looking SE at the contact on the far cliff-face the basal contact is well exposed and much higher on the cliff face than your feet on the bottom of the gorge. The same abrupt rise in the base is observed on the distal parts of certain fingers of the Maiden Creek sill. Here, the contact is concordant and bedding is actually dipping 9° to the NW. This NW dip to the bedding is consistent with the observation that the NW margin of the Trachyte Mesa laccolith is several tens of meters lower, topographically, than the SE margin. The top surface of the NE half of the Trachyte Mesa laccolith also dips to the NW. The top surface of the Trachyte Mesa laccolith in the SW has a relatively flat top. This massive sandstone layer found beneath the contact is observed wherever the basal contact is exposed. The base of the Trachyte Mesa laccolith is exposed around most of the SE margin and in a few places around the NW margin. Because the 1-m-thick sandstone layer is usually observed at the upper contact (except for Stop 1.12), and the massive sandstone layer is always found at the base, we believe the Trachyte Mesa laccolith is largely emplaced at the same stratigraphic level throughout its extent. Position yourself on the edge of the small cliff again and again look SE at the far cliff face. The top contact is also exposed on the cliff face and reveals a much thinner part of the Trachyte

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Mesa laccolith than in the gorge. This part of the Trachyte Mesa laccolith is actually a NE-oriented finger, one of four such fingers along the SE margin. We will view the termination (cross-sectional view) of this finger from Stop 1.12. Directions to Stop 1.12 Head back upstream toward Stop 1.10, but turn N up a tributary gorge after ~40 m. Head ~50 m up this gorge until you are on the top of the Trachyte Mesa laccolith. Walk ENE toward 075° for ~250 m to the cliff that marks the SE extent of the mesa. Find the only place along the cliff-face where you can scramble down several meters without scaling the cliff. You will also find 1-m-thick blocks of sandstone along the top edge of the cliff face here. This is the cliff face we saw from Stop 1.9 (on the dirt road coming in) that has sandstone between sheets of the Trachyte Mesa laccolith. Stop 1.12: Sheets, Fingers, and Intercalated Sandstone along the SE Margin of the Trachyte Mesa Laccolith GPS (UTM): 536322, 4199495. Main points: (1) This margin of the Trachyte Mesa laccolith is composed of several sheets. (2) Early sheets are full of cataclastic bands and may be a result of the emplacement of later sheets, which are undeformed. (3) The contact between sheets is marked by a thin, solid-state shear zone. (4) The sheets erode differentially here, producing an irregular erosional profile. (5) The 1-m-thick sandstone layer, usually found at the top contact, is intercalated with thin igneous sheets near the top of the mesa. (6) Looking across the valley to the SW, we see the end of the fingers that we saw from the side at Stop 1.11. These fingers are similar in shape to the ones observed at the Maiden Creek sill. Immediately SW of the sandstone blocks that are intercalated with higher level sheets, scramble down the margin several meters. Look SW at an irregular cross-sectional view of the rocks that make up the top-margin of the mesa here. Figure 11 is a photo of these rocks and outlines the borders of the sheets as they form the margin of the mesa. The middle sheet is ~1 m thick and undeformed. The top and bottom contacts are marked by a 2-cm-thick solid-state shear zone of intense foliation defined by aligned plagioclase crystals that are cataclastically crushed and elongated. These contacts between these sheets can be traced for at least 100 m to the SW. The sheets below and above are full of cataclastic bands that do not cross into the middle sheet. This brittle deformation suggests that emplacement of this middle sheet involved high strain and/or emplacement rates or that the magma already there had crystallized to a greater degree than the incoming sheet. In either case, the deformation supports forceful emplacement of the middle sheet. Clues to the mechanisms of forceful emplacement will be observed at Stop 2.1. Scramble up to the top of the cliff again and walk several meters NE to examine the 1-m-thick sandstone blocks intercalated between sheets. Note how they are undeformed and unmetamorphosed and resemble the sandstone layer on the top contact observed at Stop 1.10. This may indicate that although

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most of the laccolith was emplaced at the same stratigraphic level (just below this sandstone layer), some thin higher-level sheets were emplaced just above this sandstone layer. Look to the SW, across the valley below, at the southernmost finger that extends toward the NE (obliquely toward the viewer). Erosion reveals a cross-sectional view of this finger, which resembles two fingers side-by-side in shape (Fig. 12). Sedimentary wall rocks are immediately to the left (SE) of this “double” finger. The shape of these fingers resembles the fingers at the Maiden Creek sill. Approximately 15 m to the N of this exposure is another finger extending along the same trend (NE), but this finger is thinner. The link from this finger to the “double” finger is unclear. Directions to Stop 1.13 Walk due N for ~400 m to a small rise in the center of the Trachyte Mesa laccolith where there is a view of the NE half of the laccolith. Stop 1.13: Plateaus (Sheets?) on Top of the Trachyte Mesa Laccolith Main point: (1) The top of the Trachyte Mesa laccolith is composed of several plateaus. We suggest that these plateaus may be magma sheets. The top of the Trachyte Mesa laccolith is characterized by a series of small-scale plateaus and ridges. On the scale of the intrusion, these features are small irregularities on a relatively flat surface. On the outcrop-scale, relatively flat areas (hundreds of square meters) are often terminated by an escarpment, which rises up several meters (or less) to the next plateau or defines an elongate ridge. Ridges can be curved or straight and are continuous for several hundred meters.

Figure 11. Sheeted outcrop at Stop 1.12. The sheet outlined by the dashed line is undeformed and 1 m thick. The surrounding sheets are cut across by abundant brittle faults. A 1–2 cm thick, solid-state shear zone marks the boundary between sheets.

Looking to the NE from Stop 1.13, we can identify two main plateaus that shape the top of the NE half of the mesa. Looking to the ENE, the first plateau is slightly lower in elevation and ends at the SE margin of the mesa. There is very little vegetation and abundant exposure of flat igneous rock. This surface is slightly dipping to the NW. Moving to the NNE, a small (1–2 m) cliff defines a slightly higher, second plateau. Both of these plateaus can be traced to the NE, where they become more narrow and then further to the NE more expansive again. We suggest that these plateaus may represent individual magma sheets because their dimensions are sheet-like and they terminate along steep margins, similar to the terminations of known sheets (with wall rocks) at Maiden Creek sill. Some plateaus even terminate in bulbous margins. End of First Day Walk SE back toward Stop 1.10 and return to vehicles. Drive ~0.5 mi NE on dirt road toward Hwy 276. Turn left on Hwy 276 and drive N 5.8 mi to Hwy 95. Turn left on Hwy 95 and drive 26 mi NNW back to Hanksville. Day 2: Trachyte Mesa Laccolith (continued), Black Mesa Bysmalith, and Sawtooth Ridge Intrusions In the morning we will continue examining the Trachyte Mesa laccolith with a detailed look at the NW margin and in the afternoon examine the SE margin of the Black Mesa bysmalith and an overlook of the Sawtooth Ridge intrusion. Directions to Stop 2.1 We are driving to the same location as yesterday afternoon. Starting in Hanksville, drive 26 mi S on Utah Hwy 95. Turn right

Figure 12. Photo of two magma fingers, side-by-side at Stop 1.10 taken from Stop 1.12.

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths onto Utah Hwy 276 toward Ticaboo and reset the odometer. At mile 5.7, turn right onto a dirt road and drive ~0.5 mi. Park at a clearing by large juniper tree where the road becomes less traveled. Walk in a 350° direction for ~500 m toward the NW margin of the mesa. At the edge of the mesa you will find a basin that opens to NW, and you will have a cliff below you. Avoid the cliff in front of you and carefully make your way down, bearing to the SW until you are ~1/3 of the way down, and you will have an overview, looking to NE, of a thick (6 m) continuous red sandstone bed, which rises up onto the top of the intrusion. You should be standing on igneous rock. This is a very informative but complex outcrop, and there are several stops here. The basal and top contacts of the Henry Mountains intrusions are often observable, but this is one of the few locations in the Henry Mountains (and the only location on the Trachyte Mesa laccolith) where the wall rocks are preserved, well-exposed, and easily accessible as they are rotated upward to become the roof of the intrusion. There is also evidence for sheeting at this location and evidence for the relative timing of sheet emplacement and associated deformation.

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Stop 2.1a: Overview of Lateral Termination GPS (UTM): 536029, 4199604. Main points: (1) Geometry of the lateral termination of the intrusion. (2) There are sill-like sheets and tongue-like sheets exposed here. The edge of the laccolith is superbly exposed here. At all other margins of the Trachyte Mesa laccolith, we assume the edge of the mesa is close to the actual termination of the laccolith, but this is only an assumption based on the shape of the margin (in many places the termination is bulbous). The thickness of the Trachyte Mesa laccolith is 43 m here, which is probably close to the maximum thickness. The valley floor is at the approximate level of the base of the intrusion. The top of the intrusion is composed of at least two, and possibly three sill-like sheets here (Fig. 13). We differentiate between sill-like sheets, which extend for tens of meters in two dimensions and are a maximum of three meters thick, versus tongue-like sheets, which resemble a tongue in shape and are only exposed at the base of the laccolith here. Looking NE, sill-like sheets extend from the top of the mesa (in the SE) to beyond the edge of the mesa and rotate downward and form the actual NW margin of the intrusion.

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Figure 13. Photo and line drawing of Stop 2.1b and 2.1c at the NW margin of Trachyte Mesa laccolith where wall rocks are preserved as they rotate upward and onto the top of the laccolith. At A, there are three stacked sill-like sheets exposed. At B, tongue-like sheets are exposed. Igneous rocks are shaded.

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Most of the top sheet abruptly terminates ~1/4 of the way down the margin. The top sheet is ~50 cm thick and the termination is not tapered, but wall-like (the front surface is perpendicular to the top surface of the sheet). We assume the shape of this sheet is not erosional, but actually represents the true shape of the sheet because the thin (2–3 cm) cataclastic carapace is preserved in many places along the top and wall-like frontal face. Further down the margin, the middle sheet (below the top sheet) continues until ~1/3 the way down to the valley floor and the slope of the margin becomes steeper. This sheet is irregularly eroded and is <1 m thick. The bottom sheet is also thin (<1 m) and extends slightly further down the slope. The bottom sheet may be part of the overlying (middle) sheet. In the middle third of the slope, there are sandstone layers concordant to the top of the lowest sheet with a dip-slope of ~65°. These sedimentary beds continue to the valley floor as a dip-slope. Note that the white colored sandstone layer is full of cm-scale holes and larger cavities similar to the sandstone layer at Stop 1.10, on the SE margin, although the sandstone layer is thinner here, possibly as a result of strain. This sandstone layer comprises the inner part of the contact zone. The bottom third of this margin is composed of the sandstone beds discussed above, but they are sharply cut by tonguelike igneous sheets. These tongue-like sheets are subhorizontal and are exposed for a maximum of several meters as they protrude outward from the sandstone dip-slope. The frontal margins of these sheets are bulbous and some form perfect hemispheres whereas some are more rectangular. Below these tongue-like sheets the sandstone layers continue to dip at high angles (60–70°), but they are intensely deformed by multiple faults and currently erode into cm-scale blocks and slivers. Thin selvages of these cataclastically deformed sandstones are found on the frontal margins of some tongue-like sheets where they are even more intensely faulted and sheared. Twenty to 30 m to the NE, a ~6-m-thick massive red sandstone layer can be seen as it rises up from a relatively flat-lying position in the NW to become the top of the intrusion in the SE (Fig. 13). This sandstone bed rises up over the igneous sheets discussed above and comprises the outer part of the contact zone. Below this massive sandstone layer is a 3-m-thick series of thinner sandstone and shale beds. At the base of this 3-m-thick series is a 50-cm-thick sandstone bed that marks the inner contact zone. The actual contact with the intrusion is intensely sheared and faulted, and there are bedding plane faults between most or all of the different sedimentary layers here. The intrusion is also intensely sheared at the contact here and locally pieces of sandstone and porphyry are mixed and sheared into a fine grained, intensely foliated rock. It is important to note that the sandstone beds at the immediate contact with the laccolith, in the middle portion of the slope, have a much higher dip than the massive red sandstone does on the outer part of the contact. We suggest this attests to the resistance to bending of the thicker, and possibly stronger, massive sandstone bed and leads to a space problem between the two

layers. This space problem results in the emplacement of the tongue-like sheets and will be discussed later. The entire hillside you are standing on, which defines the edge of the mesa here, is a composite of many subhorizontal silllike igneous sheets. Most of the igneous rock (and thus most of the internal contacts) is covered by a manganese oxide coating. The irregular shape of the slope often defines the sheets. Look for areas of exposure along ridges and bumps that might be contacts. These sill-like sheets can be traced for many meters laterally and seem to about one m thick with a maximum thickness of 2–3 m. Contacts between sheets are 2–3 cm wide and are defined by a strong foliation containing highly elongated and shattered plagioclase phenocrysts that are parallel to the subhorizontal sheets. Some sheets are deformed by abundant cataclastic bands; some sheets are undeformed. These observations are similar to the deformation observed surrounding the sheet at Stop 1.12. There does not seem to be any order to the sequence in which the sheets have intruded. Based on the assumption that undeformed sheets are later than deformed sheets, later sheets intrude at various levels within the stack. In summary, an entire cross section of the contact is revealed here. We are standing on the most inner portion, where sheets are stacked. As we move 20–30 m to the NE, the sedimentary layering at the inner contact is exposed and has been rotated to high angles by the stacking of the sheets. Moving another 10–20 m NW, the outer part of the contact is exposed where a thicker sandstone layer is rotated upward and onto the top of the intrusion. Directions to Stop 2.1b Continue to carefully hike down the slope to the valley floor. Walk to the NE to examine the tongue-like sheets protruding through the sandstones at the base of the cliff discussed above. Stop 2.1b: Late Stage Magma Tongues Main points: (1) Sheets at the valley floor are tongue-like in shape. (2) These tongue-like sheets intrude through inclined (not flat-lying) sedimentary rocks and indicate that these sheets are late. (3) Sheets have bulbous shapes to their margins. These tongue-like sheets protrude through and deform the inclined sandstones that form the margin of the laccolith here (Fig. 13). This indicates that these sedimentary layers had already been rotated to their presently inclined position prior to being intruded. Because these inclined sandstones mark the lateral termination of the laccolith, these tongue-like sheets are viewed as being very late in the emplacement of the Trachyte Mesa laccolith (they postdate the formation of the lateral termination). In contrast, the sill-like sheets on top are concordant to the sandstone that represents the margin, and therefore we believe those sill-like sheets represent the earliest emplacement of magma at this margin. The sedimentary rocks below the tongue-like sheets are intensely faulted and the lateral margins of some of these sheets are also marked by faults. The bulbous margin of some of these tongue-like sheets indicates they are exposed very close to their outer contact.

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths Directions to Stop 2.1c Walk 20–30 m to the NE toward the massive red sandstone. Hike up the talus-covered slope between the exposed sheets to the SW and the red sandstone to the NE until you can walk onto the upper igneous sheets. You are almost at the top of the mesa here. Stop 2.1c: Early Sheet Intrusion and Deformation of Overlying Sedimentary Rocks GPS (UTM): 536075, 4199631. Main points: (1) There is intense solid-state deformation of the upper 2–3 cm of the silllike sheets and magmatic textures beneath. (2) The sill-like sheet intrusions are early and tilted and their slickensided surfaces exhibit evidence for late slip by subsequent intrusion. (3) There is intense brecciation of the sandstone bed immediately at the contact. (4) Flexural slip occurs on bedding planes and igneoussedimentary contacts. (5) Thinning of the massive red sandstone bed is accommodated on the grain-scale (grain size reduction and porosity collapse). Below the slickensided polish, the uppermost contact of these sheets exhibits a 2–3-cm-thick zone where the plagioclase crystals are intensely fractured, and in places completely shattered into micron-scale pieces, and these pieces were dragged along the foliation. On a macroscopic scale, the very fine grain size of the plagioclase pieces defining the foliation can be easily misinterpreted as ductile deformation of the plagioclase grains. On the surface, these pieces of plagioclase are dragged to form the obvious lineation that trends NW, perpendicular to the margin here. Shattered plagioclase grains have also been flattened to form a foliation, but the lineation is much stronger than the foliation, similar to the magmatic fabric found throughout the main body of the intrusion. The top surfaces of the sill-like sheets on the margin here are composed of thin (<1 cm), green (chlorite?), very fine grained, striated, polished surfaces forming the contact between the sheets and the sedimentary cover. This low-grade slickensided surface suggests that these upper sheets were emplaced early in the emplacement of the Trachyte Mesa laccolith, and were subsequently uplifted, rotated, and sheared during flexural slip along contacts by later emplacement of the main body of the intrusion below. Supporting this hypothesis is the observation that the sedimentary layers below these sheets are also rotated into the same inclined position, presumably by the igneous mass behind them. The other possibility is that these sheets intruded late, but followed bedding planes. In this case, as the bedding rotates downward at the margin, sheets followed bedding and intruded downward as they encountered the already formed margin. The sedimentary layers above the contact are deformed to varying degrees. Immediately above the contact, a 50-cm-thick sandstone layer is faulted and locally intensely brecciated. Some of this deformation might be related to the early emplacement of the top sheet, which terminates abruptly at this location, or it might be due to the later emplacement and inflation of the main igneous mass. There are faults between most (or all) of the sedimentary layers between the contact and the massive red sandstone

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layer, suggesting that flexural slip has occurred to partly accommodate the upward rise and rotation of these layers. Immediately at the contact, the intrusion is also intensely sheared, but only for the top few cm. Approximately one-third the distance down the margin there is a zone of mixed sedimentary and igneous rock (top sheet) that is intensely sheared. Shear bands that cut across the foliation indicate that the sedimentary rocks were moving up and over the intrusion to the SE. Detailed two-dimensional strain analysis (normalized Fry analyses) was completed on 32 sandstone samples taken from massive red sandstone layer (Fig. 14), which is a high-porosity (>10%) sandstone. Strain ratios generally increase toward the middle and upper portions of the massive red sandstone layer and then slightly decrease at the top. Two-dimensional porosity data was also collected on most of the same samples. Porosity data was collected from photographs of grains by using a graphics program whereby the area of the pore spaces versus the area of the grains could be differentiated. The thinning of the massive red sandstone is a result of grain fracturing and grain sliding, which induce a porosity decrease. In Figure 14, there is a correlation between the increase in strain ratios and the decrease in porosity along the massive red sandstone layer. Microstructurally, there is a qualitative increase in fractures as porosity decreases and strain ratios increase. Fractures mostly emanate from grain contacts and in the highly attenuated part some grains are completely crushed. Crushed grains are not observed from samples at the NW end of the layer. We associate all of this deformation to emplacement of sheets and subsequent vertical inflation of the laccolith. The observation that the strain decreases as the sandstone rolls over on to the top of the intrusion is inconsistent with the “rolling hinge” model of Hunt (1953). In this model (one of three Hunt proposed for emplacing laccoliths) an already thickened intrusion advances laterally as sedimentary rocks are rolled up and over the front of the advancing massive sheet. Return to Stop 2.1a: Evolution of a Lateral Termination Figure 15 is our interpretative cross section of the margin at Stop 2.1. We envision an incremental emplacement model for the Trachyte Mesa laccolith whereby vertical growth occurs through stacking of magma sheets. All the sill-like sheets stopped their lateral migration at generally the same location, similar to our observations at the Maiden Creek sill. There is one outcrop of a thin sheet located ~130 m N of the base of the mesa here, and this is why we place the lowermost sheet on our cross section. There does not seem to be any order to the vertical sequence of the sheets, except that the earliest sheets have been lifted to the top of the pile (the sill-like sheets that “drape” over the margin). The vertical stacking of sheets causes the mechanically strong massive red sandstone at the marginal contact to bend upward and results in a low-pressure triangle-shaped zone at the base of and in front of the marginal contact. The low-pressure zone is created because the more massive sandstone layer is thick and strong and resists bending, and therefore does not conform exactly to the

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subvertical outward margin of the growing/stacking laccolith. The thinner bedded sandstones and shales immediately at the contact are more readily deformed and conform to the outermost shape of the contact. Because the magma is under high pressure, the decrease in pressure at these voids is immediately filled by tongue-like sheets, which at this time originate from the exterior margin of the accumulating stack. In our model, these tongue-like sheets were emplaced solely as a result of this low-pressure zone. This implies that changes in magma pressure are communicated throughout the growing body, even though the body as a whole is constructed of individual sheets. This idea of magma pressure communication is supported by the observation that all the tongue-like sheets at the base arrested at the same outward distance from the Trachyte Mesa laccolith, even through they were midway through deforming upturned sandstone layers. Once these tongues filled the low-pressure zone, it once again becomes more favorable to create sheets on top or elsewhere. Therefore in our model, the location and type of sheet is partly controlled by the strength, position, and orientation of the wall rocks, which are

continually changing as more sheets are emplaced. The accumulation of sheets results in a flat-topped laccolith, which is actually the most common shape for the top of a laccolith (Corry, 1988). Directions to Stop 2.2 From the edge of the mesa here, walk toward 150° for 200 m until you see an ~1-m-thick sandstone bed eroding into blocks on top of the laccolith. Stop 2.2: Sandstone Roof of the Trachyte Mesa Laccolith Main points: (1) The current exposure of the top of the Trachyte Mesa laccolith is at or very close to the actual upper contact. (2) The same ~1-m-thick sandstone bed can be found at the upper contact on the SE, top, and NW margins, and bits and pieces of it can also be found throughout the top in the NE half of the mesa. This sandstone bed here resembles the sandstone observed at Stop 1.10, Stop 1.12, and Stop 2.1a and is lying at the contact with the igneous rock of the Trachyte Mesa laccolith below.

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths

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Directions to Stop 2.3 Walk 100 m WSW to small knob of igneous rock. Looking NE there is an overlook of the SW half of the Trachyte Mesa laccolith and looking SW there is a plateau of alluvium that extends to the SW. Stop 2.3: Overview of Internal Fabric, Geophysical Evidence for Hidden Conduits and Final Emplacement of the Trachyte Mesa Laccolith GPS (UTM): 536132, 4199471. Main points: (1) Magmatic fabric characterizes most of the intrusion. (2) The Trachyte Mesa laccolith can be divided into two zones based on orientation of magnetic lineations. The dominant orientation is to the NE, parallel to the long axis of the intrusion, and is found mainly in a linear zone running down the axis of the intrusion. (3) Magnetic anomaly data supports a NE-oriented pipe-like body at depth under the alluvium to the SW. This body is along a line extending from Mount Hillers toward Trachyte Mesa laccolith and possibly supplied magma to the intrusion. (4) Parallelism between the orientation of the (a) magnetic anomaly, (b) magnetic lineations, (c) orientation of the long axis of the intrusion, and a line that passes between Mount Hillers and the Trachyte Mesa laccolith strongly suggests that Mount Hillers is the feeder that supplied magma to the Trachyte Mesa laccolith. This is a good vantage point to discuss the fabric of the Trachyte Mesa laccolith and the AMS pattern collected from on top of the intrusion. Microstructures from the surface of this outcrop support the interpretation of a magmatic (versus solid state) fabric. Phenocrysts of undeformed, euhedral plagioclase are mostly not in contact with other crystals. The lineation, which is much

stronger than the foliation (L > S), is defined by elongate hornblende phenocrysts. In sections perpendicular to both foliation and lineation, it is often difficult to define the foliation. The AMS (magnetic) lineations from the top of the Trachyte Mesa laccolith are very shallowly plunging and vary in orientation from SE to NE with the strongest concentration to the NE (Fig. 16). Magnetic foliations are subhorizontal. Based on patterns, we have divided the lineations from the Trachyte Mesa laccolith into two domains. The NE half of the intrusion contains a central linear zone where the lineations are consistently parallel to the long axis of the intrusion (trending NE). Away from this central zone, lineations fan and diverge outward to the NNW and ESE. In the SW half of the intrusion, the pattern of lineations is more complex but seems to define a central zone that curves but is approximately E-W. South of this belt, the lineations fan outward to the S. Figure 16 also shows possible flow paths based on the magnetic lineations. The significance of these paths will be discussed later. Magnetic anomaly data was collected along three traverses over part of the alluvial plateau to the SW of the Trachyte Mesa laccolith (Fig. 17). Although the data is limited, all three traverses indicate that there is at least one large anomaly at depth that can be traced from the SW to the NE, toward the Trachyte Mesa laccolith. The anomaly peak decreases in width and increases in intensity to the NE, suggesting that the higher susceptibility material causing the anomaly is shallowing toward the NE, toward the Trachyte Mesa laccolith. A line that connects the anomalies from the three traverses (Fig. 17) broadly intersects the Trachyte Mesa laccolith in the location that the magnetic lineation pattern emanates from (Figs. 9 and 16). Nugent et al.,

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Lineation

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Figure 16. Magnetic lineations on top of Trachyte Mesa laccolith. Shaded arrows are interpreted flow paths of magma from the uppermost sheets

(2003) modeled the magnetic anomaly data (in two dimensions, using GM-SYS software) and found that they could not replicate the anomalies using dikes cutting up through sandstone. Nugent et al. (2003) were able to reproduce similar anomaly patterns using polygonal shapes of low aspect ratios at a depth of ~40 ft. There is a small outcrop of diorite porphyry located approximately halfway between Mount Hillers and the Trachyte Mesa laccolith, described (with cross sections) by Hunt (1953). The igneous body is linear in shape and aligned along a line connecting Mount Hillers to the Trachyte Mesa laccolith. One lateral margin of this body is exposed along a 2-m-high cliff with an irregular step-like contact. The actual base of the intrusion is not exposed. The other side of this body, where exposed, is very irregular in shape. The top is flat. The elevation of this body is much higher than the elevation of the Trachyte Mesa laccolith, but the position between Mount Hillers and the satellite Trachyte Mesa laccolith, as well as its alignment suggests a connection between Mount Hillers and the laccolith.

The magnetic lineation map from the top of the Trachyte Mesa laccolith (Fig. 16) suggests that the magma from the top sheets flowed along the long axis of the intrusion and spread radially outward to both sides. This pattern is particularly apparent in the NE half of the intrusion. In the SW half, the pattern of magma flow is more complex. We suggest this radial pattern is consistent with a linear centralized source that feeds sheets or fingers to both sides, but originates from the SW and is flowing generally to the NE. The magnetic lineations are also subparallel to the long axis of the plateaus in the NE half of the intrusion (Fig. 16) and support the hypothesis that the plateaus represent sill-like sheets. The map (Fig. 9) of the Trachyte Mesa laccolith illustrates that finger-like shapes protrude outward from the main body parallel to the long axis and also perpendicular to it, similar to what is observed at the Maiden Creek sill and consistent with the flow pattern obtained from the AMS lineations. We do not believe that these fingers are merely erosional remnants because

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths (a) the thicknesses of some are at a very different level than the main intrusion, and (b) some of their lateral margins bulge, similar to the true terminal margins (where wall rocks are exposed on the margins) of sheets observed elsewhere. The NE tip of the Trachyte Mesa laccolith is also shaped into two small “pinchers,” similar to how fingers seem to “pinch” at the NE margin of the Maiden Creek sill (Fig. 4). The dominant lineation orientation is along a line that can be traced directly back to Mount Hillers, 12 km away, which is also parallel to the long axis of the intrusion. The magnetic anomaly data is consistent with a magma tube (finger?) originating from Mount Hillers, which intersects the Trachyte Mesa laccolith at the location the magnetic lineations emanate from. Dikes are rare in the Henry Mountains, and dikes have not been observed anywhere near these satellite intrusions. We do not understand why magma coalesced to produce an intrusion at this location, but the data does support a model whereby intrusions are fed by magma fingers. The Trachyte Mesa laccolith may have originated as one long finger that began to branch outward into smaller fingers. Fingers expand their margins and coalesce with other fingers, which may be what we are observing at the NE margin of the Trachyte Mesa laccolith and Maiden Creek sill, where two fingers crystallized before completely coalescing. Fingers therefore evolve into sheets that are m-scale in thickness. Once a sheet is constructed, more sheets follow, but their sequence into the stack of sheets is not orderly, and partly based on wall-rock geometry and deformation. Part II: Sawtooth Ridge Intrusion and Black Mesa Bysmalith Directions to Stop 2.4 Walk SSE back to the cars and drive back NE down the dirt road until intersection with Utah Hwy 276. Turn right (S) and set odometer to zero. We are returning to the same location where we parked yesterday when we examined the Maiden Creek intrusion (see directions for Stop 1.2). Figure 18 shows the location of the stops on the Sawtooth Ridge and Black Mesa intrusions. After parking the car, walk uphill toward 225° (SW) for ~400 m. Once a reasonably flat plateau is reached, the tallest knob provides a good view of the Sawtooth Ridge and Black Mesa intrusions. Stop 2.4: Sawtooth Ridge Intrusion and Black Mesa Bysmalith Overview GPS (UTM): 535387, 4195607. Main points: (1) Sawtooth intrusion exhibits evidence for multiple intrusive sheets with bulbous terminations. (2) Black Mesa is a bysmalith (faultbounded intrusion) with flat-lying sediments on top. (3) There are weathering differences between the top ~40 m of the Black Mesa bysmalith and the main part of intrusion (~200 m). We suggest these differences may reflect the sheeted nature of the top. (4) The magma conduit to the Black Mesa bysmalith must be from bottom of the intrusion in order to raise the sedimentary “roof” rocks.

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Look WSW at the cliff face and a cross section view of the Sawtooth Ridge intrusion (Fig. 19). The Sawtooth Ridge intrusion is a highly elongate, narrow body with a jagged top from which the intrusion gets it name. The elongate direction of Sawtooth Ridge points radially away from the central part of the Mount Hillers intrusive center. The cliff face exposes the complex shape of this intrusion (Fig. 19). From this vantage point, the general direction of magma flow during emplacement of this portion of the Sawtooth Ridge was presumably ENE, or out of the cliff and toward the viewer (assuming the magma flowed parallel to the long axis of the intrusion). Bedding in the sedimentary wall rocks is bent and faulted over the top of the intrusion. Two ~5 m thick sheets, each of which ends in a bulbous termination, are clearly visible extending out from a more massive central body. We interpret these lateral “extensions” as separate sheets, based on our observations of other satellite intrusions in the region. This is our only examination of the Sawtooth Ridge intrusion. Looking WNW, the Black Mesa bysmalith is a cylindrical pluton 1.7 km in diameter and roughly 250 m thick (Fig. 20).

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Figure 19. Photo of Sawtooth Ridge intrusion taken from Stop 2.4. Cliff is ~30 m high. ig—igneous; sed—sedimentary.

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths Differential erosion between the surrounding sedimentary rocks and the igneous rocks has produced a very distinct mesa that outlines the intrusion (Fig. 21). Technically, the Black Mesa bysmalith is transitional between two forms of intrusions. The E side, which you are looking at, is a bysmalith: a piston-like, cylindrical intrusion that accommodated upward wall rock displacement mainly through faulting. Bysmaliths typically have flat roofs that are sharply truncated by vertical, curved faults (Figs. 20 and 21). The western portion of Black Mesa bysmalith is a laccolith, as the wall rocks are folded into a syncline but not faulted (Fig. 22). From this vantage point, one can observe two different erosional patterns on the lateral margin of the Black Mesa bysmalith. The top (~40 m) of the intrusion has a layered look, while the remainder of the pluton looks massive. As we will discuss at Stop 2.7, these differences in weathering coincide with distinctive patterns of AMS lineations.

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Figure 20. Photo of SSE margin of Black Mesa bysmalith taken from Stop 2.4. Cliff is ~130 m high.

Directions to Stop 2.5 Walk downhill ~300 m toward 355° and into the stream valley that runs E-W on the SE side of Black Mesa. Stop in the stream valley. Stop 2.5: Bottom of Black Mesa Bysmalith Main point: (1) Sedimentary rocks below intrusion are undeformed. The Black Mesa bysmalith is surrounded by subhorizontal strata, which abruptly change orientation and become subvertical at the contact. The roof of the bysmalith is flat, slightly N-dipping, and covered by concordant sedimentary strata of the Morrison formation. The presence of this formation at the base and the top of the intrusion in a flat lying geometry led Hunt (1953) and

Jackson and Pollard (1988) to conclude that the floor of the intrusion is close to the current bottom exposure. Based on the stratigraphic section compiled by Jackson and Pollard (1988), who estimated thicknesses through regional correlation, a maximum of 2.5 km of sedimentary rocks overlay the Morrison formation at the time of emplacement, which constitutes the lithostatic load over the roof of Black Mesa bysmalith. At this stop, we will examine the sedimentary strata of the Morrison formation immediately below the intrusion. The rocks here are essentially undeformed and unmetamorphosed.

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Figure 21. Cross section through Black Mesa bysmalith. After Habert et al. (2005).

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Figure 22. Map of Black Mesa bysmalith showing foliations and lineations. AMS-IA—anisotropy of magnetic susceptibility assisted by image analysis. After Habert et al. (2005).

Directions to Stop 2.6 Follow the stream valley uphill and to the W for ~0.5 km. Walk due N out of the streambed and uphill for ~200 m toward the cliff that delineates the southern margin of the Black Mesa bysmalith. Stop 2.6: Cataclastic Bands on the Margin of the Black Mesa Bysmalith Main point: (1) Subhorizontal cataclastic bands on the margin of the Black Mesa bysmalith are similar to those observed at the Trachyte Mesa laccolith, where they delineate separate magma sheets. On the cliff face of the Black Mesa bysmalith here, there are 1–2-m-thick subhorizontal zones rich in anastomozing cataclastic bands separated by 10–20 m of undeformed igneous rock (Habert and de Saint Blanquat, 2004). On the margins of the Trachyte Mesa laccolith, sheets are partly identified by similar zonation of brittle deformation and no deformation. Although, sheets are also delineated at the Trachyte Mesa laccolith by 2–3-cm-thick shear zones at the contacts, and this is not observed here.

Directions to Stop 2.7 Walk W ~0.5 km, contouring around the cliff and up to the saddle on the SSE margin of the Black Mesa bysmalith. Stop 2.7: Black Mesa Bysmalith in the Saddle Main points: (1) Composition and fabric of the Black Mesa bysmalith is very similar to the Maiden Creek sill and the Trachyte Mesa laccolith. (2) Complex relation (more intrusions) with main Mount Hillers body. (3) AMS lineation data indicates shallow plunges on top and moderate to steep plunges on side, which corresponds to geomorphic distinction between top and bottom of intrusion. A typical specimen of the Black Mesa bysmalith is very similar to the Maiden Creek sill or the Trachyte Mesa laccolith. It is a diorite porphyry with a microgranular porphyritic texture and consists of ~50% phenocrysts (oligoclase 30%, hornblende 10%, augite 5%, magnetite and titanite <5%, apatite <1%) and ~50% groundmass (mostly plagioclase and hornblende). The magmatic fabric at Black Mesa bysmalith is ubiquitous and is defined by the preferred orientation of plagioclase and hornblende phenocrysts. The lineation is strong but the foliation is

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths sometimes difficult to recognize in the field, similar to the Maiden Creek sill and the Trachyte Mesa laccolith. Fabric orientation varies with vertical position in the body. Subhorizontal foliations occur at the top and the bottom of the intrusion, and steeper, inward-dipping foliations occur in the middle of the intrusion. The preferred orientation of hornblende is the more obvious field structure and defines an easily measurable lineation. Ductile and/or cataclastic deformation is observed in a few locations near the top of the pluton-wall rock contact, where the first few centimeters of the intrusion are sheared. Along the margins at vertical contacts, undeformed hornblende phenocrysts are aligned parallel to vertical fault striations, suggesting that fault movement was synmagmatic. Shear sense indicators show that magma moved upward relative to wall rocks during the emplacement, which is in agreement with the tilting of the surrounding wall rocks. This fabric is characterized, on the scale of the intrusion, through the use of AMS. The magnetic foliations on the surface of the pluton (Figs. 21 and 22) define an outward dipping domeshaped pattern, which is concordant with the pluton’s margins. Thus, the pluton margin exerted a strong influence on flow and/or deformation within the interior of the evolving pluton. Steep foliations, that are sometimes inward dipping, are found along the eastern margin of the pluton and could be related to vertical flow due to peripheral fault activity. The magnetic lineations are much more instructive. Magnetic lineations within the pluton are characterized by a trimodal distribution of trends and dip (Fig. 22; also see Habert et al., 2005): (1) lineations localized on the very top of the intrusion have a general WNW-ESE trend; (2) lineations from sites below the roof of the Black Mesa bysmalith have a NNE-SSW trend; and (3) lineations localized along the eastern margin and within late dikes are vertical. Lineations are generally subhorizontal, except along the eastern margin where subvertical lineations were measured, and also in WNW-striking vertical dikes, which cut the pluton’s roof. These patterns support the inference, based on morphology, that the top of the pluton was intruded in a different manner than the lower, massive part of the intrusion. Located at the SSW end of the Black Mesa bysmalith, and in apparent morphological continuity with it, is a diorite ridge or saddle (Figs. 18 and 22). Hunt (1953) suggested this was a lateral injection zone that supplied the intrusion with magma. Our data does not support this interpretation. First, detailed mapping indicates that these two bodies are separated by the host rock of the Morrison Formation, a result that is supported by a SW-NE gravity profile (G. Habert, personal commun., 2003). Second, the two bodies have different phenocryst contents and are of slightly different petrographic facies. Last, the presence of a lateral feeder zone should have induced a symmetry in the internal fabric pointing toward the feeder, presumably toward the SSW, which is not observed in the AMS patterns (e.g., the lineation pattern is not radial from where the diorite body would connect to the body). Our observation of an axial symmetry of the fabric pattern supports a model where magma is injected from a conduit situ-

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ated below the Black Mesa bysmalith, in an approximately axial position. This is consistent with an earlier suggestion by Pollard and Johnson (1973) for a vertical feeder dyke below the bysmalith. The exact nature, geometry, and orientation of this feeder are not possible to determine precisely with our data. There are, however, the two planes of symmetry of the fabric (NNE-SSW or WNW-ESE) that may provide some constraints on the orientation of the feeder zone. Directions to Stop 2.8 Walk ~200 m N to any of several outcrops of Morrison Formation sandstone. Stop 2.8: Flat-Lying Sedimentary Rocks on Top of the Black Mesa Bysmalith Main point: (1) Flat-lying sedimentary rocks of the Morrison Formation on the top of the Black Mesa bysmalith require an upward translation of ~250 m without rotation or distortion. The roof of the Black Mesa bysmalith is flat, slightly N-dipping, and covered by concordant sedimentary strata of the lower part of the Morrison formation. These strata are equivalent to the flat lying, unmetamorphosed strata surrounding the base of the intrusion. This geometry requires that these wall rocks were translated vertically by ~250 m, without any rigid-body rotation or internal distortion. The asymmetric form of the intrusion, with its gently N-dipping roof (Figs. 21 and 22) is associated with a thickness increase toward the E, and can be attributed to the combination of (1) a primary asymmetry developed during emplacement, with more magma injected at the eastern part of the pluton, and (2) a tilt toward the NE due to the growth of Mount Hillers. Although it is difficult to observe at this location on the top of the intrusion, the fabric of the intrusion immediately below the roof rocks typically shows a pattern similar to that observed in both the Maiden Creek and Trachyte Mesa intrusions. On the western margin, a very clear stretching lineation trending 310° is measurable in the field, immediately underneath the contact, which is different from the 010° measurement obtained by AMS on a sample taken 20 cm below this contact. Given that the lineation below this layer differs for the remainder of the intrusion, we hypothesize that the topmost part of the intrusion was intruded first, probably as a sill. Directions to Stop 2.9 Walk toward 080 (ENE) for ~1 km. Please stop when you reach a precipitous drop-off. Stop 2.9: Overview and Emplacement of the Black Mesa Bysmalith Main points: (1) Comparison of evidence for sheeting in the Black Mesa bysmalith with the Maiden Creek sill and the Trachyte Mesa laccolith. (2) Emplacement model for the Black Mesa bysmalith. This location provides an overview of the three principal bodies that we have investigated on this field trip: The Maiden

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Creek sill, the Trachyte Mesa laccolith, and the Black Mesa bysmalith. We envision that each of these bodies represents a temporal succession of the three stages of pluton assembly: (1) sill intrusion, (2) inflation through stacking of sheets and accommodated by overburden bending, and (3) inflation accommodated by roof lifting along peripheral faults. The Maiden Creek sill records evidence only of the first episode, the Trachyte Mesa laccolith records the first two stages, and the Black Mesa bysmalith records evidence for all three. The clear recognition of individual magma injections may be lost in the third stage. We hypothesize, however, that this mechanism may continue to function during the later stages of emplacement. Given that the emplacement history of Maiden Creek sill and the Trachyte Mesa laccolith were covered earlier, we focus primarily on Black Mesa bysmalith. In summary, the sill-like part of the evolution of the intrusion is characterized by the WNWESE cataclastic lineations on the top of the pluton, and also the cataclastic banding we saw at Stop 2.6. These early cataclastic fabrics are also the key to the underlying magmatic fabrics. First, there is a very sharp transition from cataclastic microstructure within the first centimeters of the upper contact to the magmatic texture below. Second, this transition correlates with a change in lineation orientation from WNW-ESE at the contact to NNESSW below the contact. This latter effect was not observed in the Maiden Creek sill or the Trachyte Mesa laccolith. We present two possible models to explain the change in lineation orientation from the top to the rest of the intrusion. In both models, the lineation records the first relative movement between the magma and the wall rocks during the formation of a sill between the bottom of the Morrison Formation and the top of the Summerville Formation. In the first model, the pattern of cataclastic lineations indicates that the sill fabric was controlled by the orientation of the magma source. If we assume that the cataclastic microstructures directly record magma flow (WNW), then the underlying lineations reflect a stretching perpendicular to and away from a NNE-oriented dike. Thus, the original feeder may have been a dike (Pollard and Johnson, 1973) oriented vertically and NNE-SSW. Due to the strong temperature contrast between magma and host rocks at that time, the external part of the first pulse was chilled and cataclastically deformed during magma injection. In contrast, later magmatic fabrics within the evolving sill are controlled by the orientation of the sill plane, i.e., parallel to the gently NNE-dipping sediments, and therefore magma flowed NNE-SSW along this plane. We interpret the fabric to record strain, specifically stretching due to pluton inflation. In the second model, the lineations record motion away from a central, subhorizontal feeder, similar to the lineation pattern from the sheets on top of the Trachyte Mesa laccolith, which we suggest reflect outward, radial growth from a central, NNE-oriented zone. Here at Black Mesa bysmalith, the centralized NNE zone is covered by the sedimentary rocks on the top and we only see the margins where radial spreading is exposed. Except for the first few centimeters to meters, the only major internal fabric changes between the upper part of the pluton (the

initial sill) and the igneous mass below are the zones of anastomosing cataclastic bands separating 10–20-m-thick layers of undeformed diorite observed at Stop 2.6. The similarity between these and the zones of cataclastic bands that define sheets in the Trachyte Mesa laccolith suggest that the Black Mesa bysmalith was constructed from a series of sheets. The lack of a well-defined shear zone at the boundary between sheets, like those observed at the Trachyte Mesa laccolith, suggests that the thermal regime was different between the two intrusions (see Habert and de Saint Blanquat, 2004). It should be noted that the internal portions of the Trachyte Mesa laccolith also do not preserve internal contacts (Stop 1.11). It may be that the sheeting mechanism is only viable at the outermost contacts, or that sheeting within the hotter interiors does not result in solid-state deformation. In the sheeting growth model, the final size of the laccolith is partially controlled by the area of the initial sill. Upward bending of wall rocks occurs only at the tip of the intruding pulse. The marginal wall rocks are dominantly translated with only minor thinning because of the small area increase associated with upward growth and stretching around the intrusion. The parallelism between magmatic and solid-state lineations in the wall rocks suggests that the fabric formed during emplacement and is not late. Our proposed kinematic model is that each pulse of magma intruded subhorizontally. We do not know if later sheets intrude below or above the previous one, but based on the observations on top of the Trachyte Mesa laccolith, we assume that during the laccolith stage younger sheets are sometimes emplaced on top of older sheets. We hypothesize that the transition from Trachyte Mesa laccolith–style intrusion to the Black Mesa bysmalith–style may occur when enough sheets have accumulated so that the marginal wall rocks have been rotated to nearly vertical. At this point in the evolution of the magma chamber, faulting becomes a viable mechanism to accommodate vertical growth, partly because of the increase in area that the marginal wall rocks are subjected to during vertical accretion, and partly because the wall rocks have rotated to their maximum amount (vertical) and no further vertical growth can be accommodated by further rotation. At this point in the evolution of the intrusion, there is little evidence that sheeting continued to be the emplacement mechanism. Further, the vertical offset of the Morrison Formation to become the “roof,” indicates that the Black Mesa bysmalith was fed, at least during the major growth stage, from below. The Black Mesa bysmalith is a complex intrusion in that the margins are not similar all along the pluton’s periphery. We suggest that at least part of the laccolith stage may be preserved on the western margin. The margin along the eastern half is marked by peripheral faults, which have accommodated most of the roof lifting, whereas the margin along the western half is marked by a syncline whereby wall rocks are contiguous from the surrounding area up to the top of the intrusion. This intrusion is therefore more complex than previously described by Hunt (1953) and Pollard and Johnson (1973), and a simple plug model is not appropriate.

Sheet-like emplacement of satellite laccoliths, sills, and bysmaliths In summary, the Black Mesa bysmalith appears to have gone through a sill phase, a laccolith phase, and then a finally a bysmalith phase along its eastern margin. There is evidence for the first two phases within the Black Mesa bysmalith, but there is no definitive evidence for sheeting during the final bysmalith phase. However, the relatively large volume of magma present during inflation of Black Mesa bysmalith would cool relatively slowly and be unlikely to preserve evidence of contacts between sheets (Habert and de Saint Blanquat, 2004). These satellite intrusions only occur on the ENE side of Mount Hillers (Fig. 1). On the E, SE, and S sides of Mount Hillers, there are no satellite intrusions but the sedimentary section on the margins of Mount Hillers is rotated to steep dips (Jackson and Pollard, 1988). On the ENE side of Mount Hillers, sedimentary rocks with shallow to moderate dips can be traced high up the mountainside. The transition between these two regions of different dips occurs at the Sawtooth Ridge. We suggest that the Sawtooth Ridge intrusion may be emplaced along a scissor-fault that accommodated the differences in dips between the two regions. We also speculate that the reason that there are no satellite intrusions on the S side of Mount Hillers is because the strata were already rotated to steep dips, and therefore magma pushing outward was unable to pass easily through the steep “wall” of upturned strata, whereas on the ENE side of Mount Hillers, magma was able to flow outward along bedding planes to feed the satellite intrusions. This model assumes that the satellite intrusions were fed by sills or magma fingers, which is supported by the magnetic anomaly data SW of the Trachyte Mesa laccolith. This model also assumes that the Mount Hillers intrusive center was emplaced earlier than the satellite intrusions. End of Trip Retrace your steps. Head back north and east on the dirt road until the intersection with Utah Hwy 276. Turn left and drive north on Hwy 276, then turn left on Hwy 95 and drive north to Hanksville. ACKNOWLEDGMENTS We thank K. Charkoudian, R. Clark, G. Gleizes, A. Nugent, B. Shade, J. Silverman, and A. Stanik for assistance in the field. Dave Dilloway and the Bureau of Land Management office in Hanksville provided valuable logistical support. Funding was provided by National Science Foundation grant EAR-0003574, Centre National de la Recherche Scientifique/National Science Foundation grant 12971, and a grant from Central Michigan University to Morgan.

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REFERENCES CITED Cañón-Tapia, E., and Coe, R., 2002, Rock magnetic evidence of inflation of a flood basalt lava flow: Bulletin of Volcanology, v. 64, p. 289–302, doi: 10.1007/s00445-002-0203-8. Corry, C.E., 1988, Laccoliths; Mechanics of emplacement and growth: Geological Society of America Special Paper 220, 110 p. Engel, C.G., 1959, Igneous rocks and constituent hornblendes of the Henry Mountains Utah: Geological Society of America Bulletin, v. 70, p. 951–980. Fillmore, R., 2000, The parks, monuments, and wildlands of southern Utah: a geologic history with road logs of highways and major backroads: Salt Lake City, University of Utah Press, 268 p. Gilbert, G.K., 1877, Report on the geology of the Henry Mountains: U.S. Geographical and Geological Survey, Rocky Mountains Region, 160 p. Habert, G., and de Saint Blanquat, M., 2004, Rate of construction of the Black Mesa bysmalith, Henry Mountains, Utah, in Breitkreuz, C. and Petford, N., eds., Physical Geology of High-Level Magmatic Systems: Geological Society [London] Special Publication 234, p. 163–173. Habert, G., de Saint Blanquat, M., Horsman, E., Morgan, S., and Tikoff, B., 2005, Mechanisms and rates of non-tectonically assisted magma emplacement in the upper crust: The Black Mesa bysmalith, Henry Mountains, Utah: Tectonophysics (in press). Horsman, E., Tikoff, B., and Morgan, S.S., 2005, Emplacement-related fabric in a sill and multiple sheets in the Maiden Creek sill, Henry Mountains, Utah: Journal of Structural Geology, v. 27, p. 1426–1444. Hunt, C.B., 1953, Geology and geography of the Henry Mountains region, Utah: U.S. Geological Survey Professional Paper 228, 234 p. Hunt, C.B., 1988, Geology of the Henry Mountains, Utah, as recorded in the notebooks of G.K. Gilbert, 1875–76: Geological Society of America Memoir 167, 229 p. Jackson, M.D., and Pollard, D.D., 1988, The laccolith-stock controversy: New results from the southern Henry Mountains, Utah: Geological Society of America Bulletin, v. 100, p. 117–139, doi: 10.1130/00167606(1988)100<0117:TLSCNR>2.3.CO;2. Johnson, A.M., and Pollard, D.D., 1973, Mechanics of growth of some laccolithic intrusions in the Henry Mountains, Utah: I: Tectonophysics, v. 18, p. 261–309, doi: 10.1016/0040-1951(73)90050-4. Kelsey, M.R., 1990, Hiking and exploring Utah’s Henry Mountains and Robbers Roost: Provo, Utah, Press Publishing, 224 p. Marsh, Bruce, 2004, A magmatic mush column Rosetta Stone; the McMurdo dry valleys of Antarctica: Eos (Transactions, American Geophysical Union), v. 85, no. 47. Nelson, S.T., Davidson, J.P., and Sullivan, K.R., 1992, New age determinations of central Colorado Plateau laccoliths, Utah: Recognizing disturbed K-Ar systematics and re-evaluating tectonomagmatic relationships: Geological Society of America Bulletin, v. 104, p. 1547–1560, doi: 10.1130/00167606(1992)104<1547:NADOCC>2.3.CO;2. Nugent, A.T., Morgan, S.S., Boyd, B., Saint-Blanquat (de), M., and Horsman, E., 2003, Magnetic anomaly based interpretation of the magma feeders to the Trachyte Mesa laccolith, Henry Mts., Utah: Geological Society of America Abstracts with Programs, v. 35, no. 6, p. 339. Peterson, F., Ryder, R.T., and Law, B.E., 1980, Stratigraphy, sedimentology and regional relationships of the Cretaceous System in the Henry Mountains region, Utah, in Picard, M.D., ed., Henry Mountains symposium: Utah Geological Association Publication 8, p. 151–170. Pollard, D.D., and Johnson, A.M., 1973, Mechanics of growth of some laccolith intrusions in the Henry mountains, Utah, II; bending and failure of overburden layers and sill formation: Tectonophysics, v. 18, p. 311–354, doi: 10.1016/0040-1951(73)90051-6. Pollard, D.D., Muller, O.H., and Dockstader, D.R., 1975, The form and growth of fingered sheet intrusions: Geological Society of America Bulletin, v. 86, p. 351–363, doi: 10.1130/0016-7606(1975)86<351:TFAGOF>2.0.CO;2. Stokes, W.L., 1988, Geology of Utah: Salt Lake City, Utah Museum of Natural History and Utah Geological and Mineral Survey, 280 p.

Printed in the USA

Geological Society of America Field Guide 6 2005

Folds, fabrics, and kinematic criteria in rheomorphic ignimbrites of the Snake River Plain, Idaho: Insights into emplacement and flow Graham D.M. Andrews Michael J. Branney Department of Geology, University of Leicester, Leicester LE1 7RH, UK

ABSTRACT Recent structural analysis of the Grey’s Landing ignimbrite offers new insights into the emplacement of rheomorphic ignimbrites. We present several key localities, where volcanological and structural features reveal the emplacement history of a lavalike ignimbrite and the evolution of ductile deformation structures during and after deposition across complex topography. Excellent three-dimensional exposure allows us to interpret structural features of the Grey’s Landing ignimbrite in the context of diverse emplacement models for rheomorphic ignimbrites elsewhere and to consider field criteria to distinguish between lava-like ignimbrites and extensive silicic lavas. Keywords: rheomorphic tuff, ignimbrite, shear zone, rhyolite, sheath fold. OBJECTIVES This excursion is intended to stimulate discussion about ductile deformation processes in rheomorphic ignimbrites. The Grey’s Landing ignimbrite (Idaho, United States) provides an unrivalled opportunity to examine abundant rheomorphic structures and fabrics that reveal something of its emplacement and deformation history. We will discuss field criteria that may be used to discern a variety of emplacement and deformation models. This trip is aimed toward both physical volcanologists and structural geologists with interests in rheology, ductile deformation, and the flow of materials such as lavas, glaciers, and crustal shear zones. INTRODUCTION Ignimbrites and Rheomorphism: A History of Ideas Ignimbrites (or ash-flow tuffs) are deposits of pyroclastic density currents (Branney and Kokelaar, 2002). They typically contain various proportions of pumice and lithic lapilli in a poorly sorted matrix of ash shards. In some hot eruptions,

the pyroclasts weld together and, if sufficiently hot, the welded deposit may flow as a coherent ductile mass, a process known as rheomorphism. Rheomorphic ignimbrites typically contain a penetrative welding fabric, flow folds, and an elongation lineation, such as stretched vesicles and prolate fiamme (Fig. 1). They typically exhibit microscopic to 10-m-scale folds and, in some cases, upper autobreccias. A variety of kinematic indicators include an oblique foliation (sometimes described as imbrication of fiamme), rotated porphyroclasts, boudinaged fiamme, and inclined tension cracks (e.g., Schmincke and Swanson, 1967; Wolff and Wright, 1981; Branney and Kokelaar, 1992). Various models have been proposed for the emplacement and ductile deformation of rheomorphic ignimbrite (Fig. 2); these are not mutually exclusive, and different rheomorphic ignimbrites may have different deformation histories. Current research is aimed at constraining (1) what causes the rheomorphic deformation, (2) its timing relative to deposition, and (3) the duration and style of deformation, as these differ between the existing models. Welding in ignimbrites has commonly been inferred to postdate emplacement so that the vertical welding profile reflects

Andrews, G.D.M., and Branney, M.J., 2005, Folds, fabrics, and kinematic criteria in rheomorphic ignimbrites of the Snake River Plain, Idaho: Insights into emplacement and flow, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 311–327, doi: 10.1130/2005.fld006(15). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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Figure 1. Schematic block diagram showing structures formed by rheomorphism in ignimbrites of the Mogàn Formation, Gran Canaria (after Schmincke and Swanson, 1967).

Figure 2. Schematic diagrams summarizing the emplacement of a rheomorphic ignimbrite as inferred from the four most readily accepted models. From Sumner and Branney (2002).

Insights into emplacement and flow the thermal and compaction profiles through the ignimbrite after emplacement (Ross and Smith, 1961). However, Schmincke and Swanson (1967) interpreted rheomorphic deformation in peralkaline ignimbrites to occur while a hot pyroclastic current decelerates, gradually deflates, agglutinates, and comes to a halt en masse (Fig. 2A). In this conceptual model, the welding and rheomorphism starts prior to, and independently of, any postemplacement cooling and welding. Chapin and Lowell (1979) invoked progressive agglutination, accretion and shear within a laminar boundary layer at the base and sides of a valley-filling plug-like pyroclastic current that behaved as a Bingham fluid (Fig. 2B), forming L1 and F1 structures, before slumping in a hot state toward the valley axis (perpendicular to the earlier flow direction) producing F2 folds (Fig. 3). In contrast, Wolff and Wright (1981) proposed that welding and rheomorphism begin only after the pyroclastic density current has ceased transport and deposition (Fig. 2C); the stationary deposit welds in situ due to internal heat and loading and then is able to remobilize and flow down-slope en masse. Branney and Kokelaar (1992) proposed that ignimbrites deposit by rapid, progressive aggradation from the base of a sustained pyroclastic density current and that in some cases welding and rheomorphism may start during deposition (Figs. 2D and 4); in this scenario, rheomorphic deformation may initially be partitioned in a subhorizontal ductile shear zone near the top of the deposit, and this migrates upward as the deposit continues to aggrade (Fig. 4). The rheomorphism may then continue after the pyroclastic current has dissipated (t3 of Fig. 2D). An implication of this model is that one cannot assume the ignimbrite sheet was initially isother-

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Figure 3. Schematic block diagram showing two-phase rheomorphism as inferred by Chapin and Lowell (1979) from the Wall Mountain Tuff, Colorado. L1—elongation lineation formed by stretched pumice lapilli and vesicles.

mal, and some ignimbrite welding profiles may reflect in part the differing rheologies of successive pyroclast populations supplied by the current to the site of deposition with time. Sheath folds (Fig. 5) with axes subparallel to elongation lineations have been discovered in many rheomorphic ignimbrites (Branney et al., 2004). This discovery means that the two-phase deformation of Chapin and Lowell (1979) and Schmincke (1990) is more readily interpreted as recording a single sustained, progressive deformation event. Ideas thus have paralleled those

Figure 4. Schematic diagram showing the development of rheomorphic structures in a syndepositional shear zone at the current-deposit interface during progressive aggradation of ignimbrite (modified from Branney and Kokelaar, 1992).

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Figure 5. Schematic block diagram showing the development of a sheath fold during progressive simple shear.

developed for tectonic shear zones. Critical evidence in support of this is that most ignimbrites have only one elongation lineation and that in some (e.g., the Grey’s Landing ignimbrite), the azimuth orientation of this lineation and the sheath fold axes change with height in the ignimbrite, which may reflect changing shear directions during deposition. Terminology Welding in Tuffs Welding is the adhesion, plastic deformation, and compaction of hot pyroclasts (Ross and Smith, 1961). According to pyroclast rheology, welding may vary from rapid agglutination and even coalescence during deposition to slower loading-compaction during cooling of a thick ignimbrite sheet (e.g., Freundt, 1998). High- and Extremely High-Grade Ignimbrite Grade refers to the intensity of welding exhibited by ignimbrite sheets (Walker, 1983; Branney and Kokelaar, 1992). The grade continuum may be divided into four intergradational categories: (1) Low-grade ignimbrites are predominantly nonwelded and may be emplaced at temperatures down to ambient. (2) Moderate-grade ignimbrites have both welded and nonwelded zones with little intense welding and generally no rheomorphism. (3) High-grade ignimbrites are predominantly welded, with intensely welded zones, and commonly exhibit local rheomorphism. (4) Extremely high-grade ignimbrites are intensely welded even to their upper surfaces. They typically exhibit rheomorphic structures and include lava-like lithofacies (e.g., the Grey’s Landing ignimbrite).

Agglutination and Coalescence Agglutination is the very rapid welding of depositing hot pyroclasts prior to burial compaction (e.g., in some basaltic spatter deposits). Coalescence is the homogenization of hot pyroclasts back to a coherent viscous fluid on deposition, so that the original clast outlines disappear (Branney and Kokelaar, 1992) (e.g., in the Grey’s Landing ignimbrite, and in some clastogenic lavas [Furukawa and Kamata, 2004]). Lava-Like “Lava-like” is a purely descriptive term referring to a lithofacies that resembles a lava in that it lacks visible vitroclastic textures. It may be massive or flow-banded. This term can be used for parts of extremely high-grade ignimbrites, where pyroclasts are inferred to have coalesced. It does not imply any particular origin or emplacement mechanism. Rheomorphism Rheomorphism is the ductile deformation of hot, welded pyroclastic material during and/or just after deposition; this is separate and distinct from tectonic deformation. Ignimbrites undergoing rheomorphism develop a variety of ductile deformation structures, including flow-banding, flow-folds, folded and attenuated pumices, vesicles, and welding fabrics (Schmincke and Swanson, 1967; Wolff and Wright, 1981, Branney et al., 2004). These features develop in the viscous welded mass while it is still hot and degassing. For rheomorphism to occur, pyroclasts must be sufficiently fluidal at the time of welding to readily deform. This condition is favored by high emplacement temperatures, high eruptive mass-flux, minimal ingestion of atmospheric

Insights into emplacement and flow air into the density current during transport, rapid deposition, strongly peralkaline chemistries, and/or high dissolved volatile (e.g., H2O, Cl, F) contents (Mahood, 1984). Progressive Aggradation Progressive aggradation is the incremental increase in the thickness of a deposit with time because of deposition. Fisher (1966) and Branney and Kokelaar (1992, 1997) proposed that deposits of most pyroclastic density currents, including massive layers in ignimbrites, deposit incrementally beneath a sustained current (Fig. 4). This is in contrast to the conceptual model in which a pyroclastic density current progressively deflates during runout causing it to decelerate and ultimately halt en masse. Rates of aggradation may vary (unsteadiness), and this may impart layering within an ignimbrite. Sheath Fold “Sheath fold” is a purely descriptive term for strongly curvilinear folds where the measured fold hinge is parallel or subparallel to the stretching lineation. Sheath folds are interpreted to be produced by high finite, inhomogeneous strain. They characteristically appear as eye-structures when viewed parallel to the stretching direction (Fig. 5) and commonly occur associated with oblique folds. Branney et al. (2004) have shown that sheath folds and oblique folds are common in rheomorphic ignimbrites from diverse volcanic settings (e.g., Gran Canaria, Colorado, Pantelleria). Themes for Discussion throughout the Field Trip Is There Any Link between Rheomorphism and Pyroclastic Emplacement? Is rheomorphism affected by pyroclastic transport and deposition, or is it entirely “secondary” and thus independent of the emplacement mechanisms? For example, if rheomorphic shear accompanies transport and deposition (Fig. 2), evidence of these processes may be “frozen” into the ignimbrite and be used to provide insights into pyroclastic emplacement processes. Rates of Welding and Rheomorphism Welding has traditionally been considered to occur on time scales of weeks to years (e.g., Riehle et al., 1995). However, high- and extremely high-grade ignimbrites are now commonly interpreted to undergo syndepositional and rapid welding in just seconds to minutes (Mahood, 1984; Branney and Kokelaar, 1992; Freundt, 1998). Welding rate is principally controlled by the pyroclast viscosity and the exerted stress. It may be that ignimbrites interpreted to have undergone load-welding were cooler (e.g., ≤750 °C) than those interpreted to have undergone agglutination and rheomorphism (e.g., 750–1050 °C). What evidence can be used to constrain rates of welding and rheomorphic deformation? The Deformation History How can the deformation history be unraveled from rheomorphic structures and fabrics? Was the deformation polyphase

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and/or progressive? Was the strain partitioned spatially, and did this change with time? Distinguishing Extremely High-Grade Ignimbrites from Lavas Given the very different eruption and emplacement mechanisms (e.g., explosive versus extrusive), what criteria can be used to decide whether a unit is a lava-like ignimbrite or a true lava? Criteria proposed include the presence or absence of remnant pyroclastic layering, vitroclastic textures, autobreccias, crystal breakage, lithic concentrations, and overall morphology with respect to substrate topography (e.g., Ekren et al., 1984; Bonnichsen and Kauffmann, 1987; Branney et al., 1992; Henry and Wolff, 1992; Sumner and Branney, 2002). Some of these criteria are not reliable and some can be obscured by deformation and devitrification. As with lavas, extremely high-grade ignimbrites tend to exhibit scarce lithic clasts and lower crystal breakage relative to lower grade ignimbrites (Branney and Kokelaar, 1992). Conversely, lavas commonly develop local vitroclastic textures (fiamme and welded shards) by hot shearing of pumiceous autobreccia (Pichler, 1981; Manley, 1995). Amongst the best criteria are the presence of widespread basal autobreccia (Henry and Wolff, 1992), only known to occur in lavas, and topographydraping veneer-like, thin (<10 m), sheet-like morphologies with tapering margins, which are characteristic of ignimbrites. However, such features may not be exposed, so can structural analysis be used? Lavas can, for example, inherit folds and fabrics from conduit walls, whereas ignimbrites must develop their folds and fabrics develop after eruption. Similarities and Differences to Tectonic Shear Zones The structures in the Grey’s Landing ignimbrite are similar to those commonly found in exhumed ductile shear zones. Both typically are characterized by strong, noncoaxial plane strain, and have a strong, subhorizontal L = S fabric with intrafolial folds, and sheath folds. Cleavage, however, does not develop in rheomorphic ignimbrites, which are largely noncrystalline at the time of deformation. This can hinder structural analysis. Tectonic shear zones are confined—bound on either side by nondeforming rock mass—whereas rheomorphic ignimbrites have a free upper surface. Tectonic shear zones tend not to migrate, whereas rheomorphic shear zones may migrate during aggradation. Crustal shear zones often intersect layered sequences (e.g., bedded sediments) with an inherent mechanical anisotropy, which strongly influences the deformation (e.g., a thick quartzite will deform differently from a sequences of interbedded, thin shales and sandstones). In contrast, the Grey’s Landing ignimbrite may have been almost mechanically isotropic during deformation, and the deformation is more akin to deformation within glaciers and lavas. REGIONAL SETTING AND STRATIGRAPHY The Yellowstone–Snake River Plain volcanic province (Fig. 6) is dominated by large-volume (>10 km3) rhyolite lavas and ignimbrites, erupted from a series of major eruptive centers

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that show an age progression from McDermitt caldera (ca. 16 Ma) in the southwest to the Yellowstone Plateau (ca. 2–0.6 Ma) in the northeast (Pierce and Morgan, 1992). The systematic age progression appears to be related to the southwestward movement of the North American plate over a fixed thermal anomaly, or hot spot. Presently, there is renewed debate as to the nature of this hot spot, with two opinions: (1) it represents a deep-rooted “mantle plume-tail” related to the mantle plume that produced the Columbia River Basalts (e.g., Pierce and Morgan, 1992); and (2) it represents an upper mantle convection anomaly at the base of the lithosphere (e.g., Humphreys et al., 2000). The Yellowstone–Snake River Plain volcanic province is one of the most important regions for the study of large volume rhyolite eruptions (Bonnichsen, 1982; Christiansen, 2001; Ekren et al., 1984; Bonnichsen and Kauffman, 1987; Manley, 1996). Canyons and escarpments across southwestern and southern Idaho expose voluminous (<1000 km3) rhyolite sheets that include long lavas and extensive ignimbrites erupted from inferred eruptive centers (Owyhee-Humboldt, Bruneau-Jarbidge, and Twin Falls, between ca. 15 and ca. 8 Ma; Fig. 6). Mid- to late-Miocene rhyolites of southern Idaho and northern Nevada have been consigned to the Idavada Group (Malde and Powers, 1962). This includes the Cougar Point Tuff Formation (Bonnichsen and Citron, 1982), thought to be derived from the Bruneau-Jarbidge eruptive center; the Cassia Mountains succession, thought to be derived from the Twin Falls eruptive center (McCurry et al., 1996); and the Rogerson Formation, which occupies the intervening region between Twin Falls, Idaho, and Jackpot, Nevada (Andrews et al., 2006; Fig. 6). The Rogerson Formation (Fig. 7) comprises seven ignimbrite members intercalated with volcaniclastic sediments and paleosols (Andrews et al., 2006). It records at least eight large explosive eruptions with intervening repose periods. The ignimbrites are predominantly fine-grained and intensely welded. Most are welded and two are rheomorphic and lava-like, although nonwelded rhyolitic pyroclastic layers also occur. They are typical Snake River Plain high-silica rhyolites (e.g., Hughes and McCurry, 2002), with sparse anhydrous plagioclase and 2pyroxene glomerocrysts and Fe-Ti oxides. No vents have been identified, and the Rogerson Formation eruptions may have been located at the Bruneau-Jarbidge and Twin Falls eruptive centers, or in the intervening area (Fig. 6). The Grey’s Landing ignimbrite Member (type locality: Grey’s Landing Recreation Ground, Idaho; Fig. 8) is a 5–65-mthick rhyolite sheet. It comprises a stratified ashfall deposit, overlain by a largely lava-like ignimbrite with a lower vitrophyre, a thick lithoidal center, and a thin upper vitrophyre, locally overlain by a nonwelded top (Figs. 9 and 10). The ignimbrite is intensely rheomorphic; it is compositionally zoned with vertical variations in glass and pyroxene crystal chemistry (Andrews et al., 2006). Eruption temperature is estimated at 950 °C–1050 °C (Andrews et al., 2006). No source vent has been identified, nor has a reliable radiometric age been confirmed; however, the ignimbrite is demonstrably younger than the 10.54 Ma Rabbit Springs ignimbrite (Bill Bonnichsen, 2004, personal commun.).

FIELD TRIP This field trip will visit the Grey’s Landing ignimbrite of the Rogerson Formation (Fig. 7) and, if time allows, the neighboring House Creek ignimbrite. We will examine a small number of key localities around the Rogerson Graben. The itinerary has been ordered so as to begin by introducing participants to the graben and its rhyolite stratigraphy before examining the range of rheomorphic structures, starting with the simplest structures and then building up to a more complex and complete picture. This requires some backtracking each day; this is not a “linear” road-log based trip. Logistics All localities are (at the time of writing) on publicly owned open land and most are easily accessible by a combination of short walks and vehicles. Most localities are accessible to twowheel drive/low-clearance vehicles provided care is taken. Relevant topographic maps: 1:100,000—Rogerson; 1:24,000—Cedar Creek, Browns Bench North, Meteor, Salmon Butte. There are several hotel-casinos in Jackpot, Nevada, together with a general store and gas station. There is a gas station and small store in Rogerson, Idaho. As with all desert fieldwork, remember to bring sufficient water and sunblock. This excursion does not require strenuous effort; however, care should be taken: local hazards include fast traffic, steep and loose talus slopes, snakes, and hunters (late August–January). Please refrain from hammering at the outcrops—the exposed surfaces are often more informative than fresh alternatives, and the rocks are visually stunning—some offer genuine aesthetic pleasure; they make much better photographs. The talus slopes contain abundant fresh, hand-specimen–sized fold closures and lineated foliation surfaces. Day 1 Introduction to the succession and to welding and fusing in tephras. Syndepositional rheomorphic folds, elongation lineations, and kinematic indicators. Depart Jackpot, Nevada, heading north on U.S. Highway 93 (U.S. 93). Stop 1.1—Backwaters: stratigraphic context, welding and fusing of tephras; Stop 1.2—Roadkill (U.S. 93): thin ignimbrite on stratified ashfall deposit; Stop 1.3—Norton Canyon Road: intermediate-thickness ignimbrite with rheomorphic folds; Stop 1.4—Grey’s Landing: thick ignimbrite with abundant small-scale sheath folds. Directions to Stop 1.1 Drive north from Jackpot, Nevada, on U.S. 93 for ~7 mi (11 km), and turn left for Backwaters Recreation Area (signposted). Continue along a gravel track to the parking area at the base of the large cliffs on the right (north).

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Figure 6. Simplified geologic map of the Yellowstone–Snake River Plain volcanic province. Modified from Pierce and Morgan (1992); Hughes and McCurry (2002); and Rodgers et al. (2002). Age data taken from Hughes and McCurry (2002).

Figure 7. General vertical section through the Rogerson Formation in the Rogerson Graben. Soils, erosion surfaces, and unconformities are marked. Details of lithofacies, welding grade, type locality, minimum volume, and crystal assemblage are given on the right.

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Figure 8. Simplified geological map of the Rogerson Graben and surrounding area (Twin Falls County, Idaho, and Elko County, Nevada), showing the present distribution of the Grey’s Landing ignimbrite (gray) and field trip stops. Selected UTM coordinates for Zone 11T shown. Adapted from Andrews et al. (2006).

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Figure 9. Vertical division into two structural zones at Grey’s Landing (stops 1.1, 1.4, 2.1, 2.2, and 2.3). The lower zone (0–20 m) is predominantly a penetrative, subhorizontal foliation with intrafolial folds and sheath folds. The upper zone (20–55 m) is dominated by meter-scale to 10-m-scale complex folding and refolding of an earlier penetrative and intrafolial fabric. We interpret the earlier fabric to be the same as that which dominates the lower zone.

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Figure 10. Graphic log of the Grey’s Landing ignimbrite in the area around Stop 1.3.

Stop 1.1: Backwaters (UTM Zone 11: 685154E 4659673N) This locality provides an introduction to the Rogerson Formation (Fig. 7) and the Rogerson Graben (Fig. 8) and is a good place to consider the geological context for the remainder of the excursion. The Rogerson Graben is a NNE-SSW–trending half-graben that deepens westward toward the Browns Bench Fault. The footwall of the fault has uplifted a thick succession of extensive rhyolitic ignimbrites that comprise the Browns Bench escarpment on the western horizon. The ignimbrites are high-grade and may derive from the Bruneau-Jarbidge eruptive

center (Bonnichsen and Citron, 1982; see Fig. 6) 30 km to the west. The Rogerson Graben is one of at least 10 similarly trending half-graben on the southern side of the Snake River Plain. They are thought to relate to Basin and Range extension, and the volcanism seems to have accompanied the extension. The Rogerson Formation (Andrews et al., 2006) occupies the Rogerson Graben and postdates the ignimbrites exposed in the escarpment. It includes seven ignimbrites and associated tephras (Fig. 7) erupted from the Snake River Plain to the north. Unconformities and variations in ignimbrite thickness

Insights into emplacement and flow indicate that eruption and deposition were contemporaneous with graben extension. The lowest exposed unit in the formation, the Jackpot rhyolite (Fig. 7), is best observed on the track in and out of the Backwaters recreation area. It forms the cliffs and rolling surface to the south. It is overlain by poorly exposed bedded volcaniclastic sediments, the Rabbit Springs ignimbrite (a vitric layer at road level), some younger, poorly exposed volcaniclastic sediments, and the Browns View ignimbrite, which forms a discontinuous ledge at base of cliffs to the east. The overlying buff-colored and poorly exposed Backwaters ignimbrite is partly fused by the overlying Grey’s Landing ignimbrite, which is the main cliffforming unit on either side of the small canyon. Examine accessible parts of the Rabbit Springs and Browns View ignimbrites on the approach to the base of the Grey’s Landing ignimbrite. The basal contact relations of the Grey’s Landing ignimbrite reveals that the massive vitrophyre overlies a stratified, crystal-rich vitrophyre that, in turn, rests on a baked paleosol (Fig. 9). There is no autobreccia at the base of the ignimbrite. Examine the underlying Backwaters ignimbrite, noting how parts are fused to black glass. Themes to consider: How do these deposits compare to less welded ignimbrites? Why are fiamme and lithic clasts absent? How might the deposits be distinguished from lavas? Does the stratified vitrophyre represent a fused ashfall deposit? Exactly where, within the vitrophyre, is the base of the Grey’s Landing ignimbrite? Why is the underlying paleosol baked rather than fused? Directions to Stop 1.2 Return to U.S. 93, noting how the Grey’s Landing ignimbrite (the topmost exposed unit) thins dramatically toward the east. Turn right (south) and drive ~1 mi (1.6 km) and stop in the large pull-off on the right (west) side of U.S. 93. Cross the road to examine the exposure in the road-cut. Take care; the exposure is named Roadkill for a reason! Stop 1.2: Roadkill (UTM Zone 11: 692234E 4660906N) Along U.S. 93, the Grey’s Landing ignimbrite is an entirely vitric sheet, ≤5 m thick. It sits on a sequence of stratified ash and pumice-rich layers that overlie a paleosol. The upper 0.5–1 m of the stratified ash is interpreted as an ash-fall deposit from the Grey’s Landing eruption (Andrews et al., 2006). It correlates with the stratified, crystal-rich vitrophyre at Stop 1.1. Examine the basal contact of the ignimbrite and note how the stratified ash deposits become increasingly fused toward the base of the ignimbrite. The ignimbrite does not have a basal autobreccia, and the underlying ash deposit is not deformed. The ignimbrite is massive glass, and neither bedding nor rheomorphic structures are discernible within it. Next, note the stretched vesicles (E-W) at the top surface of the outcrop (best seen along the very edge). Orientated thin sections from this locality show that the entire ignimbrite is intensely sheared, top to the west, even at the base. Themes to consider: What criteria distinguish between lava-like ignimbrites and lavas? What is the significance of the

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absence of autobreccia? Can rhyolite lavas be emplaced as very thin sheets? Directions to Stop 1.3 Return north on U.S. 93 and turn onto the Norton Canyon road (signposted). Pull up on the right after ~250 m, on reaching the first set of low crags on the right. Here the Grey’s Landing ignimbrite is 6–10 m thick (Fig. 10) and forms the local topographic surface. It overlies the same partly fused ashfall deposit as at Stops 1.1 and 1.2 (poorly exposed here) and, again, the base of the ignimbrite lacks an autobreccia and the underlying fused ash is not deformed. Stop 1.3: Norton Canyon Road (UTM Zone 11: 690615E 4663972N) This locality introduces the effects of devitrification as well as some macroscopic rheomorphic structures. The ignimbrite comprises three subhorizontal and subparallel zones defined by differences in devitrification: (A) a basal vitrophyre, (B) a central lithoidal zone, and (C) an upper vitrophyre (see Fig. 10). The basal and upper vitrophyres represent the original glassy state of the unit during deposition (as at Stop 1.2). The jointed, redbrown central lithoidal zone represents the devitrified center of what was probably originally a glassy hot deposit. Devitrification is the static crystallization of volcanic glass during prolonged cooling from high initial emplacement temperature. It is associated with volume loss that encourages the formation of joints, and it can obscure vitroclastic textures. Abundant deformation structures occur within the central lithoidal zone and upper vitrophyre but are absent in the basal vitrophyre. Note, within the lithoidal zone, a strong stretching lineation, curvilinear flow folds (picked out by the pervasive jointing), rotated porphyroclasts (probably crystals), and sheared and oblique vesicles. Meter-scale, recumbent isoclines on the top surface of the upper vitrophyre have hinge lines parallel to the vesicle stretching direction (E-W). Structural data collected at this and adjacent localities are presented in Figure 10; note the dominant subhorizontal foliation and intrafolial nature of the fold axial planes and the girdle-like distribution of fold hinges parallel to oblique to the lineation. Themes to consider: What is the structural style? How does this compare to styles seen at tectonic shear zones and/or mylonite zones? What of folds on the upper surface? How much of the deposit has undergone rheomorphism? Is there any evidence that mechanical layering has influenced the rheomorphism? Can existing emplacement models account for structures such as these? Given how thin the deposit is, how long could deposition and/or deformation have continued? What role, if any, might devitrification exert on deformation? Why are the clearest structures found in the devitrified material? Directions to Stop 1.4 Return to U.S. 93 and turn left (north), then turn left (west) for Grey’s Landing Recreation Ground (signposted). Proceed to Grey’s Landing and park near the slipway.

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Stop 1.4: Grey’s Landing (UTM Zone 11: 687822E 4666778N) First walk along the lake shore at the base of the northern cliffs and examine the base of the Grey’s Landing ignimbrite; take care crossing the extensive talus slopes. Note that the stratified ashfall tuff at the base is totally fused here, and the underlying paleosol is baked (similar to Stop 1.1; Fig. 9). Fossil grass imprints are preserved on the top surface of the paleosol. The basal part of the ignimbrite is massive vitrophyre that lacks macroscopic rheomorphic structures such as folds and autobreccia. Examine the base of the much thicker, central lithoidal zone (cf. Stop 1.3); a “devitrification front” is made up of spherulites that protrude downward into the vitrophyre. Where within the basal vitrophyre is the contact between ignimbrite and fall deposit? What is the origin of the boundary between vitric and devitrified rhyolite? Walk south along the base of the cliffs (beware of loose talus) and onto the prominent bench developed in devitrified ignimbrite. The ignimbrite immediately below the bench is thoroughly penetrated by closely spaced, anastomosing joints and exhibits no macroscopic folds. The lowest macroscopic folds, including excellent sheath folds, occur at the top of the bench, where the jointing is more hackley and traces out abundant intrafolial fold closures on vertical surfaces. Subhorizontal foliation surfaces have a strong lineation parallel to fold hinges. Many surfaces exhibit scallopshaped “dimple joints” on which the lineation is absent. Stretching lineations and fold hinges are consistently parallel and trend E-W at this level within the ignimbrite (lower stereonet Fig. 9). We shall see that with increasing height in the ignimbrite, this orientation gradually changes, first in a counter-clockwise sense and then, above 33 m, in a clockwise sense. This has been interpreted as recording gradual changes in the rheomorphic transport direction with time, while the ignimbrite progressively aggraded (Branney et al., 2004). Themes to consider: Is the style of folding here similar to Stop 1.3? What is the origin of the diverse forms of jointing? Has the jointing been folded, or did it mimic (develop along) already folded flow-banding? When did it form relative to rheomorphism? Why does the lowermost lithoidal rhyolite lack rheomorphic structures? Summary Day 1 familiarized participants with the stratigraphy and structure of the Rogerson Graben and gave an impression of the intensity of welding and rheomorphic deformation in the Greys Landing ignimbrite, along with its devitrification history. We have examined the ubiquitous subhorizontal fabric in lower parts of the ignimbrite, and have seen that lowermost folds include intrafolial sheath folds. We have observed that at a given height in the ignimbrite, the orientation of fold axes is subparallel to the elongation lineations. Day 2 Nature of “late” rheomorphism in thick ignimbrite.

Stop 2.1—Upper part of Grey’s Landing ignimbrite: largescale folding of the earlier fabric; Stop 2.2—“Too Cool”: later, large-scale folding in 3D; Stop 2.3—Cedar Creek Reservoir: later folding at the upper surface. Directions to Stop 2.1 Depart Jackpot, Nevada, heading north on U.S. 93. Return to Grey’s Landing Recreation Ground and park at the slipway (originally Stop 1.4). Stop 2.1: Upper Parts of Grey’s Landing Ignimbrite (UTM Zone 11: 688087E 4666732N) Examine the exposures above the talus slope on the north side, immediately adjacent to and above the slipway. Ascend a narrow gulley near the corner of the cliff; climb out of the gulley onto the flat promontory on the left. Examine the curved cliff face behind the promontory, noting the complex folding of the earlier intrafolial, subhorizontal fabric that we saw yesterday. Largescale folds are not intrafolial; they are upright to recumbent and tight to isoclinal. Descend out of the gulley, and then walk east along the base of the cliffs, examining the large-scale folds and refolding of the earlier intrafolial structures. Note that the large-scale folds are curvilinear and trend ~E-W. Gradually ascend the slope, and pause where the cliffs turn north into a small side-canyon. Look east across the sidecanyon at the opposite face and appreciate the large-scale eye-structures developed near the top of the cliff. Cross the side-canyon and ascend to examine the complex refolding patterns produced around the large sheath fold. While descending the slope and returning to the slipway, note how the dominant fabric is intrafolial away from the effects of refolding by later, large-scale folds. Structural data collected within the complexly refolded zone at this locality is presented in Figure 9. Note the spread of foliation and fold axial plane data around the horizontal and the preferred trend of lineation and fold hinges E-W. This supports observations that the later folds are curvilinear to sheath-like and commonly recumbent. Themes to consider: What is the style of folding? How do these folds relate to the earlier, intrafolial and smaller folds? Does an increase in fold scale indicate a change in rheological conditions? What form has the transition between levels dominated by initial folding and those dominated by subsequent folding? Directions to Stop 2.2 Drive back out from Grey’s Landing and turn left (north) through the prominent iron gate at the junction of four tracks. Drive along a rough track ~1.5 mi (3.1 km), then descend the prominent fault scarp to the shore in a small N-S–trending cove (not suitable for 2WD or low-clearance vehicles). Stop 2.2: Too Cool (UTM Zone 11: 687268E 4668336N) Examine the low, subhorizontal surface immediately west of the beach. This surface is the eroded remnant of the original upper

Insights into emplacement and flow surface of the Grey’s Landing deposit. Different colored layers (due to variations in devitrification) are folded into a series of 2-mscale domes and basins, complex refolded folds, and sheath folds. Explore this area and, if possible, descend to the base of the low cliffs to examine these features in three dimensions. Themes to consider: What is the style of folding? How do the folds relate to structures seen at lower levels in the ignimbrite, e.g., at Stops 1.4 and 2.1? Can any effect of proximity to the original surface be discerned? Did the differently colored layers exhibit different mechanical properties during the deformation? Directions to Stop 2.3 Return to U.S. 93 and turn left (north) and then turn left (northwest) on to unnamed road immediately after milepost 13. Turn left (west) onto the Murphy Hot Springs Road and cross Salmon Dam; continue west to Cedar Creek Reservoir (signposted). Pass around the north side of the outbuildings and turn right (north) toward the slipway (signposted). Pull off the track to the left and stop on the broad, flat area overlooking the reservoir, after ~1 mi (1.6 km). Stop 2.3: Cedar Creek Reservoir (UTM Zone 11 673712E 4674456N) Walk west to where the wire fence reaches the cliff line. Then follow the cliff line toward the northeast, noting the folded internal layering within the ignimbrite (Fig. 11). This layering is produced by different devitrification and welding intensities within the Grey’s Landing ignimbrite. Although thicknesses of the layers vary, their stratigraphic order remains consistent. The uppermost layer is a nonwelded, orange tuff that will be seen again at Stop 3.4 (Salmon Dam) resting upon the upper vitrophyre. The layering is folded into upright, 10-m-scale anticlines and synclines with hinges trending NW-SE (Fig. 11). Many folds are strongly curvilinear, and axes plunge steeply to both the NW and SE. Several of the anticlinal closures are composite and include more than one generation of earlier fold closure within them. This produces refold interference patterns in which more than three generations of folds interact. Continue for ~400 m to the northeast until the cliff line peters out. If the reservoir level is sufficiently low (typically in August– November), descend to the lake bed and walk southwest along the base of the cliffs. Excellent sheath folds and refolded folds can be seen along the cliffs and in loose blocks on the lake bed. Some surfaces preserve dilational fractures of vesiculated ignimbrite around fold closures. Themes to consider: What is the style of folding? Does the scale and wavelength of the folds inform us about rheology of the unit? How are these folds produced: are they buckle folds produced by coaxial shortening, or are they similar-style flow folds produced by continued flow during emplacement and modified by gravity? Summary At the end of Day 2, participants should have learned something of the nature and style of post-emplacement rheomorphism

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as demonstrated by the complex refolding of the earlier ubiquitous subhorizontal fabric. We can draw two major conclusions today: (1) there has been a second generation of folds (phase of rheomorphism) in the Grey’s Landing ignimbrite; and (2) we can characterize this second (generation) phase of rheomorphism. The duration of this rheomorphism can be constrained using cooling models. The second phase of rheomorphism remains enigmatic. The fold axes trend east-west, similar to that of the early folds. Could the later folds have developed as buckle-style, coaxial shortening folds during a phase of downslope creep perpendicular to the initial emplacement direction, or could they simply have been generated by continued flow in the same direction as the initial emplacement? Critical to this interpretation is that there is no second set of stretching lineations perpendicular to the late fold axes and at high angles to the initial stretching lineation. Participants are encouraged to consider the alternatives and look for supporting evidence. The cooling history of the ignimbrite affects its rheological evolution, and hence, its deformation history. Viscosity within silicic glasses is inversely proportional to temperature for a given composition and volatile content. This relationship is strongly nonlinear; at the glass transition temperature (~650 °C–725 °C for anhydrous rhyolite), viscosity increases by several orders of magnitude. At low temperatures and short timescales, below the glass transition, the glass responds to strain as a brittle solid (Webb, 1997). At high temperatures and short timescales, above the glass transition, it behaves as a non-Newtonian fluid and will flow. Therefore, cooling is an important limiting factor on the duration and style of deformation. Thermal profiles can be established by adapting cooling models (e.g., Manley, 1992) for magmatic temperature estimates obtained from crystal analyses and geothermometry. This has been done for the 60-m-thick, intensely welded (low porosity) Grey’s Landing ignimbrite, emplaced at 1000 °C onto porous substrate and then buried by a thin, porous ash deposit. This model deals with conduction of heat to the air and the substrate, and does not account for convection of fluids and/or volatiles, or advection of hot material due to rheomorphic folding. It predicts rapid (<1 yr) cooling of the upper and lower parts (now vitrophyres) and more gradual cooling of the interior of the ignimbrite. By incorporating the glass transition, the duration of ductile behavior at different depths can be estimated. The maximum duration of ductile behavior in the center is ~15 yr; however, given the exclusion of heat loss by convection and advection, this is probably a considerable overestimate. Day 3 (Half Day) Stop 3.1—Monument Canyon: transition from a domain preserving dominantly effects of “early” rheomorphism to a domain recording early deformation modified by “late” rheomorphism; Stop 3.2 (optional)—House Creek: rheomorphism in an incredibly thin ignimbrite; Stop 3.3 (optional)—West Bay: intrafolial sheath folds and spaced zones of “late” rheomorphism;

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Figure 11. Simplified geologic map of Stop 2.3, Cedar Creek Reservoir, showing the location of fold axial traces within the Grey’s Landing ignimbrite.

Insights into emplacement and flow Stop 3.4 (optional)—Salmon Dam: large-scale synforms at the upper surface with infolded orange tuff. Directions to Stop 3.1 Depart Jackpot, Nevada, heading north on U.S. 93. Follow the route for Day 2 across Salmon Dam, and turn right (south) onto a gated track ~300 m east of the junction for Cedar Creek Reservoir. Proceed along this rough track (not suitable for low ground-clearance vehicles) until you reach a large water tank and corral. Stop 3.1: Monument Canyon (UTM Zone 11 675840E 4671116N) This stop involves a short walk to consider how rheomorphism was influenced by substrate topography. It provides the opportunity to draw together understanding of rheomorphic features seen in the previous two days. Walk south along the track and onto the ridge to the immediate right. Gradually walk westward, examining the low crags at the top of the ridge to look for fold closures in the central lithoidal zone. The Grey’s Landing ignimbrite in this vicinity forms a thin feather-edge. Is there any rheomorphism at this location? Has devitrification and/or jointing obscured any macro-structures? Over the next 400 m, the ignimbrite thickens to >15 m at the southern entrance to Monument Canyon. Meter-scale stretched vesicles are exposed in overhangs at head height; they parallel the local stretching direction (E-W) and give a top-to-the-west shear sense. Curvilinear, recumbent isoclines are abundant at the entrance to Monument Canyon, and the strong, subhorizontal fabric is obvious in cliffs on the opposite side. The ignimbrite continues to thicken over the next 200 m to the north before undergoing a dramatic change. The attitude and nature of the subhorizontal, intrafolial fabric changes abruptly across a steep ~E-W trending boundary (parallel to stretching direction and edge of the deposit). To the north of this boundary, the subhorizontal fabric is folded and refolded at different scales (meter scale to ten meter scale). The later folds are characteristically upright to inclined, open through to isoclinal, and plunge gently to steeply. Participants should range across the canyon flanks, making sure to look back across the canyon at the opposite flank. Themes to consider: How is the earlier, subhorizontal fabric affected by later folding? Is there evidence of a time break between early and later folding? How do these structures relate to the two alternative explanations given for the “late” rheomorphism (see Summary, Day 2)? Directions to Stop 3.2 Return to the Murphy Hot Springs Road and turn left toward Salmon Dam. Continue west ~8 mi (12 km) and pull off to the left opposite a large sand pit. Walk northeast ~250 m to a small abandoned quarry. Stop 3.2: House Creek Quarry (UTM Zone 11 666463E 4668768N) The line of low (1–2 m) crags forming the edge of the quarry is the House Creek ignimbrite (Bonnichsen et al., 1989). Note

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how thin the ignimbrite is (~5 m original thickness) and the development of sheath folds and a strong stretched vesicle lineation (Branney et al., 2004). No autobreccia has been observed in this deposit. Themes to consider: What is the structural style? Are there any kinematic indicators that can be used to determine the transport direction of ductile shear? How can rheomorphism occur within such a thin unit? What is the significance of the apparent absence of autobreccia? Directions to Stop 3.3 Drive eastward, past the turnoffs for Stops 2.3 and 3.1. Turn right (south) onto a rough track running down to a flat area on top of a low slope above West Bay. Note the rotation of fault blocks on west-dipping, listric, normal faults in the middle distance. Stop 3.3: West Bay (UTM Zone 11 686473E 4675065N) Examine the low crags and exposed surfaces at the top of the slope, gradually tracking south. Here the deposit is at least 50 m thick; subhorizontal L = S fabrics have been gently warped by open, upright folds, which are most clearly seen across West Bay in the foreground. Sheath folds are well exposed throughout this section. Walk slowly along the cliff line southward, noting that the foliation reveals narrow “steep belts” of intensely refolded deposit, spaced ~50 m apart. Themes to consider: Can the L = S tectonites (including sheath folds and strong stretching lineation) at this location be accounted for in the same way as those at Stops 1.2, 1.3, and 1.4? How does the style of later folding here relate to that seen at Stops 2.2–3.1? Directions to Stop 3.4 Return to the road and turn right; drive ~100 m and pull off onto the open ground on the left at the top of the grade. Park vehicles and walk back to the dam. Beware traffic appearing around the corner. Stop 3.4: Salmon Dam (UTM Zone 11 687018E 4675787N) To the north of the dam, the Salmon Falls Creek canyon can be seen cutting through Pliocene basalts erupted from Salmon Butte (multiple dark gray flows on right side; Bonnichsen and Godchaux, 2002) and the Grey’s Landing ignimbrite (red-brown exposures on the left). The Grey’s Landing ignimbrite is exposed below two basalt lavas where the dam meets the west side of the canyon. Observe the variation in color and structure in the steep face, from the dam and the far side of the canyon. A large, upright syncline within the upper part of the Grey’s Landing ignimbrite has preserved a partly enclosed remnant of massive, orange tuff, which varies from nonwelded to slightly fused near the contact with the ignimbrite upper vitrophyre. Take extreme care on these slopes. Close examination of the contact relations reveals that the tuff was deposited when the ignimbrite was hot, ductile, and still deforming. Please do not hammer.

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Themes to consider: Is the orange tuff an upper nonwelded part of the ignimbrite (note the Grey’s Landing ignimbrite generally lacks pumice lapilli), or is it an ashfall deposit (e.g., of co-ignimbrite origin)? What is the effect of continued deposition of nonwelded tephra? What is the nature of the contact between welded and nonwelded tephra? CONCLUDING REMARKS We leave it for individuals to decide which (if any) of the various preexisting models for rheomorphism best accounts for the deformation features in the Grey’s Landing ignimbrite. A structural analysis (authors’ work in preparation) suggests that the ubiquitous small-scale intrafolial folds and associated strong lineation initiated within in a narrow shear zone (~0.5–1 m thick) during pyroclastic deposition and agglutination. With time, this shear zone migrated upward along with the rising aggradation surface (Fig. 4) so that, by the time the entire thickness of ignimbrite had aggraded, all levels had been subjected to intense shear. The layer of orange ash from the same eruption was then deposited on top (Stop 3.4). Below this, ductile, gravity-driven shear continued, probably at decreasing strain rates but affecting an increased thickness of the ignimbrite. The larger scale of the later sheath folds and oblique folds reflects the greater thickness of the shearing layer. The steeper attitudes of some of these folds reflects the presence, during the deformation, of an upper free surface, which has no direct analogy in crustal shear zones. The later deformation affected the upper orange ash and so must have developed after pyroclastic emplacement had ceased. We infer that at this time the shear rate was decreasing as cooling and degassing caused changes in rheology. The last stages of deformation were partly brittle, with the formation of tension cracks in upper parts of the ignimbrite. Overall, the deformation history was probably progressive (rather than comprising two entirely separate fold phases), and the style of deformation evolved in response to the evolving stresses and rheologies. Analysis of the Grey’s Landing ignimbrite is continuing, in order to characterize the deformation and develop a model for the ignimbrite emplacement and rheomorphism. ACKNOWLEDGMENTS Many thanks to Bill Bonnichsen, Marty Godchaux, and Mike McCurry for introducing us to some of these wonderful localities and their continued discussion, advice, and encouragement. Curtis Manley kindly allowed us to apply his cooling model to the Grey’s Landing ignimbrite, and Craig White kindly provided equipment support. Nancy Riggs, Jocelyn McPhie, John Stix, Guillermo Labarthe-Hernandez, Gerardo Aguirre-Diaz, and Jorge Aranda-Gomez provided much-needed practice for GDMA’s viva examination on different occasions during the fieldwork. GDMA was supported by Natural Environment Research Council (NERC) studentship NER/S/A/2001/06292 and the University of Leicester.

REFERENCES CITED Andrews, G.D.M., Branney, M.J., Bonnichsen, B., and McCurry, M., 2006, Rhyolitic ignimbrites in the Rogerson Graben, southern Snake River Plain volcanic province: volcanic stratigraphy, eruption history and basin evolution: Bulletin of Volcanology (in press). Bonnichsen, B., 1982, Rhyolite lava flows in the Bruneau-Jarbidge Eruptive Center, southwestern Idaho, in Bonnichsen, B., and Breckinridge R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau Mines and Geology Bulletin, v. 26, p. 283–320. Bonnichsen, B., and Citron, G.P., 1982, The Cougar Point Tuff, southwestern Idaho, in Bonnichsen, B., and Breckinridge, R.M., eds., Cenozoic Geology of Idaho: Idaho Bureau Mines and Geology Bulletin, v. 26, p. 255–281. Bonnichsen, B., and Godchaux, M.M., 2002, Late Miocene, Pliocene, and Pleistocene geology of southwestern Idaho with emphasis on basalts in the Bruneau-Jarbidge, Twin Falls, and Western Snake River Plain regions, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and magmatic evolution of the Snake River Plain volcanic province: Idaho Geological Survey Bulletin, v. 30, p. 233–312. Bonnichsen, B., and Kauffman, D.F., 1987, Physical features of rhyolite lava flows in the Snake River Plain volcanic province, southwestern Idaho, in Fink, J.H., ed., The emplacement of silicic domes and lava flows: Geological Society of America Special Paper 212, p. 119–145. Bonnichsen, B., Christiansen, R.L., Morgan, L.A., Moye, F.J., Hackett, W.R., Leeman, W.P., Honjo, N., Jenks, M.D., and Godchaux, M.M., 1989, Excursion 4A: Silicic volcanic rocks in the Snake River Plain—Yellowstone Plateau province, in Chapin, C.E., and Zidek, J., eds., Field excursions to volcanic terranes in the western United States; Volume II, Cascades and Intermountain West: New Mexico Bureau of Mines and Mineral Resources, v. 47, p. 135–182. Branney, M.J., and Kokelaar, P., 1992, A reappraisal of ignimbrite emplacement: progressive aggradation and changes from particulate to non-particulate flow during emplacement of high-grade ignimbrite: Bulletin of Volcanology, v. 54, p. 504–520, doi: 10.1007/BF00301396. Branney, M.J., and Kokelaar, B.P., 1997, Giant bed from a sustained catastrophic density current flowing over topography: Acatlàn ignimbrite, Mexico: Geology, v. 25, p. 115–118, doi: 10.1130/0091-7613(1997)025<0115: GBFASC>2.3.CO;2. Branney, M.J., and Kokelaar, P., 2002, Pyroclastic density currents and the sedimentation of ignimbrites: London, The Geological Society Memoir 27, 143 p. Branney, M.J., Kokelaar, P., and McConnell, B.J., 1992, The Bad Step Tuff: a lava-like ignimbrite in a calc-alkaline piecemeal caldera, English Lake District: Bulletin of Volcanology, v. 54, p. 187–199. Branney, M.J., Barry, T.L., and Godchaux, M., 2004, Sheathfolds in rheomorphic ignimbrites: Bulletin of Volcanology, v. 66, p. 485–491, doi: 10.1007/s00445-003-0332-8. Chapin, C.E., and Lowell, G.R., 1979, Primary and secondary flow structures in ash-flow tuffs of the Gribbles Run paleovalley, central Colorado, in Chapin, C.E., and Elston, W.E., eds., Ash-flow tuffs: Geological Society of America Special Paper 180, p. 137–154. Christiansen, R.L., 2001, The Quaternary and Pliocene Yellowstone Plateau volcanic field of Wyoming, Idaho, and Montana: U.S. Geological Survey Professional Paper 729-G, 145 p. Ekren, E.B., McIntyre, D.H., and Bennett, E.H., 1984, High-temperature, largevolume, lavalike ash-flow tuffs without calderas in southwestern Idaho: U.S. Geological Survey Professional Paper 1272, 73 p. Fisher, R.V., 1966, Mechanism of deposition from pyroclastic flows: American Journal of Science, v. 264, p. 350–363. Freundt, A., 1998, The formation of high-grade ignimbrites; 1: Experiments on high- and low-concentration transport systems containing sticky particles: Bulletin of Volcanology, v. 59, p. 414–435, doi: 10.1007/s004450050201. Furukawa, K., and Kamata, H., 2004, Eruption and emplacement of the Yamakogawa rhyolite in central Kyushu, Japan: A model for emplacement of rhyolitic spatter: Earth Planet Space, v. 56, p. 517–524. Henry, C.D., and Wolff, J.A., 1992, Distinguishing strongly rheomorphic tuffs from extensive silicic lavas: Bulletin of Volcanology, v. 54, p. 171–186. Hughes, S.S., and McCurry, M., 2002, Bulk major and trace element evidence for a time-space evolution of Snake River Plain rhyolites, Idaho, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and magmatic evolution of the Snake River Plain Volcanic province: Idaho Geological Survey Bulletin, v. 30, p. 161–176.

Insights into emplacement and flow Humphreys, E.D., Dueker, K.G., Schutt, D.L., and Smith, R.B., 2000, Beneath Yellowstone; evaluating plume and non-plume models using Teleseismic images of the upper mantle: GSA Today, v. 10, no. 12, p. 1–7. McCurry, M., Watkins, A.M., Parker, J.L., Wright, K., and Hughes, S.S., 1996, Preliminary volcanological constraints for sources of high-grade, rheomorphic ignimbrites of the Cassia Mountains, Idaho: Implications for the evolution of the Twin Falls Volcanic Center: Northwest Geology, v. 26, p. 81–91. Mahood, G.A., 1984, Pyroclastic rocks and calderas associated with strongly peralkaline volcanic rocks: Journal of Geophysical Research, v. 89, p. 8540–8552. Malde, H.E., and Powers, H.A., 1962, Upper Cenozoic stratigraphy of the western Snake River Plain, Idaho: Geological Society of America Bulletin, v. 73, p. 1197–1210. Manley, C.R., 1992, Extended cooling and viscous flow of large, hot rhyolite lavas: implications of numerical modelling results: Journal of Volcanology and Geothermal Research, v. 53, p. 27–46, doi: 10.1016/0377-0273(92)90072-L. Manley, C.R., 1995, How voluminous rhyolite lavas mimic rheomorphic ignimbrites: eruptive style, emplacement conditions, and formation of tuff-like features: Geology, v. 23, p. 349–352, doi: 10.1130/0091-7613(1995)023<0349: HVRLMR>2.3.CO;2. Manley, C.R., 1996, Physical volcanology of a voluminous rhyolite lava flow: the Badlands lava, Owyhee plateau, SW Idaho: Journal of Volcanology and Geothermal Research, v. 71, p. 129–153, doi: 10.1016/03770273(95)00066-6. Pichler, H., 1981, Italienische Vulkan-Gebeite III: Lipari, Vulcano, Stromboli, Tyrrenisches Meer: Sammlung Geologische Fuhrer, v. 69, p. 1–233. Pierce, K.L., and Morgan, L.A., 1992, The track of the Yellowstone hotspot: volcanism, faulting, and uplift, in Link, P.K., Kuntz, M.A., and Platt, L.P., eds., Regional geology of eastern Idaho and western Wyoming: Geological Society of America Memoir 179, p. 1–53.

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Riehle, J.R., Miller, T.F., and Bailey, R.A., 1995, Cooling, degassing, and compaction of rhyolitic ash-flow tuffs: a computational model: Bulletin of Volcanology, v. 57, p. 319–336. Rodgers, D.W., Ore, H.T., Bobo, R.T., McQuarrie, N., and Zentner, N., 2002, Extension and subsidence of the Eastern Snake River Plain, Idaho, in Bonnichsen, B., White, C.M., and McCurry, M., eds., Tectonic and magmatic evolution of the Snake River Plain volcanic province: Idaho Geological Survey Bulletin, v. 30, p. 121–155. Ross, C.S., and Smith, R.L., 1961, Ash-flow tuffs: their origin, geologic relations and identification: U.S. Geological Survey Professional Paper 366, p. 1–81. Schmincke, H.-U., and Swanson, D.A., 1967, Laminar viscous flowage structures in ash-flow tuffs from Gran Canaria, Canary Islands: Journal of Geology, v. 75, p. 641–664. Schmincke, H.-U., with contributions by Freundt, A., Ferriz, H., Kobberger, G., Leat, P., 1990, Geological Field Guide: Gran Canaria: Witten, Germany, Pluto Press, 202 p. Sumner, J.M., and Branney, M.J., 2002, The emplacement history of a remarkable heterogeneous, chemically zoned, rheomorphic and locally lava-like ignimbrite: “TL” on Gran Canaria: Journal of Volcanology and Geothermal Research, v. 115, p. 109–138, doi: 10.1016/S0377-0273(01)00311-0. Walker, G.P.L., 1983, Ignimbrite types and ignimbrite problems: Journal of Volcanology and Geothermal Research, v. 17, p. 65–88, doi: 10.1016/03770273(83)90062-8. Webb, S.L., 1997, Rheology, relaxation and the glass transition in silicate melts: Reviews of Geophysics, v. 35, p. 191–218, doi: 10.1029/96RG03263. Wolff, J.A., and Wright, J.V., 1981, Rheomorphism of welded tuffs: Journal of Volcanology and Geothermal Research, v. 10, p. 13–34, doi: 10.1016/ 0377-0273(81)90052-4.

Printed in the USA

Geological Society of America Field Guide 6 2005

Mesozoic lakes of the Colorado Plateau Timothy M. Demko* Department of Geological Sciences, University of Minnesota–Duluth, Duluth, Minnesota 55812, USA Kathleen Nicoll University of Oxford, School of Geography and the Environment, Oxford OX1 3TB, UK Joseph J. Beer Department of Geological Sciences, University of Minnesota–Duluth, Duluth, Minnesota 55812, USA Stephen T. Hasiotis University of Kansas, Department of Geology and the Natural History Museum and Biodiversity Research Center, Lawrence, Kansas, 66045-7613, USA Lisa E. Park University of Akron, Department of Geology, Akron, Ohio 44325-4101, USA

ABSTRACT The Upper Triassic Chinle Formation and the Upper Jurassic Morrison Formation preserve a record of lacustrine deposition along the western margin of tropical Pangaea and post-Pangaean North America. The lake deposits in these formations contain archives of sedimentary and geochemical paleoclimatic indicators, paleoecological data, and characteristic stratal architecture that provide glimpses into the evolution of basins linked to global- and continental-scale tectonic events and processes, and the establishment of a mosaic of continental paleoecosystems. This field trip highlights the lacustrine and associated fluvial deposits of the Monitor Butte Member of the Chinle Formation and the Tidwell and Brushy Basin Members of the Morrison Formation in the southern part of the Colorado Plateau region, with emphases on: (1) sedimentary facies analysis and paleogeography of the paleolakes; (2) stratal architecture and high-frequency sequence stratigraphy; (3) recognition of lake basinfill types; and (4) paleontology and ichnology of lake strata and their paleoecologic, paleohydrological, and paleoclimatic interpretation. Keywords: lakes, Chinle, Morrison, Colorado Plateau, paleoclimate, paleoenvironments, paleoecology.

*E-mail: [email protected]. Demko, T.M., Nicoll, K., Beer, J.J., Hasiotis, S.T., and Park, L.E., 2005, Mesozoic lakes of the Colorado Plateau, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 329–356, doi: 10.1130/2005.fld006(16). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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330 INTRODUCTION

structing the paleoenvironments through a critical interval of geologic time. The Morrison Formation is known worldwide to contain a large diversity of dinosaur fossils including some of the largest herbivores that ever roamed the planet. Since the great dinosaurbone wars of the late 1800s, geologists and paleoecologists have been interested in the details of the numerous habitats of tropical western Pangaea. The most recent studies of the Morrison Formation interpret the distribution of environments and their associated biotic communities, describing the Late Jurassic extinct ecosystem as part of a complex landscape mosaic, the components of which have shifted through time (Carpenter et al., 1998; Gillette, 1999; Turner and Peterson, 2004). The importance of interpreting continental paleoecological archives associated with the Mesozoic sedimentary succession of the Colorado Plateau is underscored in studies of the early diagenetic mineralogy (Turner and Fishman, 1991), geochemical isotopes (Dunagan and Turner, 2004), preserved invertebrate fauna (Schudack et al., 1998; Good, 2004), associated flora (Parrish et al., 2004), continental ichnofossils (Hasiotis, 2004), and paleosols (Demko et al., 2004) of lacustrine and related strata. The study of continental waterways in this succession of ancient landscapes—namely, its rivers, floodplains, wetlands, and lakes—are of particular interest, because water controls

This field trip highlights some of the Mesozoic fluviopalustrine-lacustrine deposits of the Colorado Plateau region of southern Utah and examines the evolution of ancient lake systems and their associated paleoecosystems reconstructed from the sedimentary and paleontologic record. Data used in these reconstructions include stratal architecture, paleosols, continental ichnofossils, and a rich fossil fauna and flora. Primary themes of the trip include (1) Permian-Jurassic stratigraphy; (2) continental depositional systems; and (3) Pangaean to post-Pangaean tectonics, paleogeography, paleohydrology, and paleoclimate. Field stops examine fluvio-palustrine-lacustrine deposits formed under tropical monsoonal climatic conditions, including those of the Upper Triassic Chinle Formation in Glen Canyon National Recreation Area and Capitol Reef National Park, Utah (Fig. 1). Other stops examine palustrine-lacustrine deposits formed under tropical wet-dry climate conditions, including the Upper Jurassic Morrison Formation in Moab, Four Corners (Utah, Colorado, New Mexico, and Arizona), Henry Mountains, and Capitol Reef National Park areas, Utah (Fig. 1). The continental succession of Mesozoic strata of the Western Interior of the United States provides a valuable basis for recon-

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the occurrence and distribution of organisms and their activities (e.g., Parrish, 1989; Fiorillo et al., 2000; Engelmann et al., 2004; Hasiotis, 2004). Understanding the significance of aquatic deposits and their associated physical, chemical, and biological components is particularly useful in reconstructing a conceptual paleohydroclimatic framework (the ancient hydrologic and climatic setting) for a particular depositional unit. Current models of modern and ancient lacustrine depositional systems interpret strata in the context of three major facies associations at various scales from beds to members or parasequence to depositional sequence to sequence-set at the scale of meters to hundreds of meters. These facies associations exhibit a characteristic stacking pattern as lake basins fill (see Bohacs et al., 2000, and references therein). Three end-member lithofacies associations are recognized generally on objective physical, chemical, and biological criteria: fluvial-lacustrine, fluctuating profundal, and evaporative (Carroll and Bohacs, 1999). The fundamental controls on these lithofacies associations in space and time include lake morphometry and water depth. These controls are a function of the relative balance of rates of potential accommodation change (eustatically and tectonically forced) and sediment + water supply (hydroclimatically-forced) (Einsele and Hinderer, 1998). Models predicting lake occurrence, distribution, and character link the three most common facies associations with distinctive lake-basin types: overfilled, balanced-fill, and underfilled lake basins (Bohacs et al., 2000). The fluvio-lacustrine deposits examined on this field trip will be described and interpreted within this lake basin classification system.

extending from ~85°N to 90°S (Ziegler et al., 1983; Blakey et al., 1993) (Fig. 2A). Global sea level was low throughout the Permian and the Triassic (Vail et al., 1977), and Pangaea disrupted nearly every part of the zonal circulation due to its large size, resulting in a high degree of continentality of climate. The location of the large landmass in low latitudes and the presence of a warm seaway that acted as a moisture source maximized summer heating in the circum-Tethyan part of the continent (Parrish, 1993). The resultant Pangaean climate was likely seasonally wet-dry, or megamonsoonal (Parrish et al., 1986; Dubiel et al., 1991). As Pangaea broke up and North America moved north, the exposed land area was distributed more evenly on either side of the equator. Seasonality intensified through the Triassic, and the equatorial regions and mid-latitude continental interiors became more arid when the monsoonal circulation was at its maximum (Parrish, 1993). Further breakup of Pangaea eventually disrupted the megamonsoonal circulation pattern, and global climate gradients became more latitudinal. However, global climate models (GCMs) (Moore et al., 1992; Valdes and Sellwood, 1992) also suggest that a semiarid to arid climate persisted in the Western Interior through the Late Jurassic. A high-pressure system dominated southwestern North America, with surface temperatures of 30–40 °C in the summer and 0–20 °C during the winter. Estimated rainfall amounts and precipitation-evapotranspiration (P/E) ratios from the GCMs (Moore et al., 1992; Valdes and Sellwood, 1992) suggest direct meteoric contribution may have been minimal.

PALEOGEOGRAPHIC SETTING AND STRATIGRAPHY

The Chinle Formation was deposited in a broad, fully continental, cratonic basin created by subsidence due to viscous flow in the mantle associated with the subduction of the Farallon plate and flexure due to supracrustal loading by the associated volcanic

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arc (Lawton, 1994). The vast, >2.5 million km2 Chinle basin was along the tropical west coast of the supercontinent Pangaea between 5–15°N paleolatitude (Dubiel, 1994) (Fig. 2A). The Chinle Formation consists principally of fluvial, floodplain, palustrine, and lacustrine deposits with minor eolian and playa environments present at the close of Chinle time (Blakey and Gubitosa, 1984; Dubiel 1989b). Various petrographic, sedimentary isopach, stratigraphic, and paleontologic evidences suggest that the sediment sources were the Mogollon Highlands in Arizona and the Uncompahgre and Front Range highlands of the Ancestral Rocky Mountains in Colorado. The Mogollon Highlands and the adjacent magmatic arc system contributed both volcanic and sedimentary detritus to the lower part of the Chinle Formation in the southern part of the basin (Stewart et al., 1972; Riggs et al., 1996). The lower part of the Chinle Formation (Shinarump, Monitor Butte, and Moss Back Members) was deposited in a succession of valley-fill sequences under monsoonal climatic conditions (Demko, 1995; Demko et al., 1998). The upper part of the Chinle Formation (Petrified Forest, Owl Rock, and Church Rock Members) was deposited in a regionally dynamic basin complex of alluvial-lacustrine systems (Stewart et al., 1972; Dubiel, 1989a, 1994) (Fig. 3). Vertic paleosols and cyclic lacustrine facies are among the sedimentologic and paleopedologic evidence that suggest the Chinle basin was characterized by strongly seasonal precipitation, with distinctive wet and dry seasons (Dubiel et al., 1991). The Chinle Formation is well known as one of the richest Late Triassic fossil plant–bearing units in the world (Ash, 1980; Demko et al., 1998), with over 70 plant taxa, including lycopods, ferns, cycads, conifers, bennettitaleans, seed ferns, and several other unclassified forms in the published literature. The fossil floral physiognomy (Ash 1967, 1972, 1980; Ziegler et al., 1993) and vertebrate paleoecology (e.g., Parrish et al., 1986; Parrish, 1989) are among the paleontologic evidence that contribute to a reconstruction of the Chinle paleoenvironment. The Chinle Formation preserves an abundant and diverse continental ichnofauna upon which much of the paleohydrologic interpretations of Triassic tropical Pangaea have been made (Hasiotis and Dubiel, 1993, 1994, 1995a, 1995b; Hasiotis and Mitchell, 1993; Hasiotis et al., 1993, 2004). Trace-making organisms and trace fossils can be placed into behavioral categories that indicate the space, trophic associations, and groundwater moisture zones occupied by organisms (Hasiotis, 2000). The tiering of above- and below-ground trace-making organisms in Chinle deposits indicates that their distribution was controlled in part by annual and seasonal fluctuations of unsaturated (soil moisture) and saturated (water table and phreatic) zones, which in turn was controlled by regional climate (Hasiotis and Mitchell, 1993; Hasiotis and Dubiel, 1994) (Fig. 4). Traces of crayfish, bees, beetles, soil bugs, and plant roots are the preserved products of the water balance in paleosols that record the relation between annual precipitation inputs, solar radiation, evapotranspiration losses, and soil moisture. This information, combined with other paleontologic, sedimentologic, stratigraphic, isotopic, and paleogeographic data, suggests spatial and temporal variations

associated with tropical monsoonal climates during deposition of the lower Chinle Formation to increasingly arid climate at the end of Chinle deposition (e.g., Dubiel and Hasiotis, 1994a, 1994b, 1995). For example, the great depth, wide distribution, and high abundance of crayfish burrows in the Shinarump, Temple Mountain, Petrified Forest, and Owl Rock Members suggest that influent rivers were fed by the local and regional saturated zone (Hasiotis and Mitchell, 1993; Hasiotis et al., 1993). Adhesive meniscate burrows (AMB), constructed by beetles (adults and larvae) or soil bugs, co-occur with rhizoliths and the upper parts of crayfish burrows, reinforcing the interpretation of moderate soil moisture levels in the unsaturated zone (Hasiotis and Dubiel, 1994, 1995b). Upper Jurassic Morrison Formation This unit, famous for abundant dinosaur fossils, but also containing abundant and diverse plant, invertebrate, and trace fossils, was deposited throughout the Rocky Mountain region from New Mexico to Montana between 30 and 45°N paleolatitude (Peterson, 1994; Chure et al., 1998) (Fig. 2B). The Morrison Formation represents 7–8 m.y. of deposition from latest Oxfordian or early Kimmeridgian (ca. 155 Ma) to early Tithonian (ca. 148 Ma) (Kowallis et al., 1998; Litwin et al., 1998). The Morrison Formation includes the Tidwell, Salt Wash, and Brushy Basin Members in the Colorado Plateau area. The Tidwell Member interfingers with the Bluff Sandstone and Junction Creek Sandstone Members in the Four Corners region, whereas the lower Brushy Basin and Salt Wash Members grade into and interfinger with the Recapture and Westwater Canyon Members in the same area (Peterson and Turner-Peterson, 1987; Peterson, 1994) (Fig. 3). The Morrison Formation consists of a succession of conglomerate, sandstone, siltstone, mudstone, limestone, and evaporites that were deposited in alluvial, lacustrine, palustrine, eolian, and continental-marine transitional environments (Brady, 1969; Peterson and Turner-Peterson, 1987; O’Sullivan, 1992; Peterson, 1994; Dunagan, 1998; Turner and Peterson, 1999). Many of the alluvial, lacustrine, palustrine, and eolian deposits within the Morrison Formation were modified by some degree of pedogenesis after deposition, producing a variety of immature to mature paleosols, some of which mark significant unconformities and can be correlated across the region (Demko et al., 2004). Paleoclimatic interpretations for the Morrison ecosystem range from tropical wet-dry to arid, depending on the various indicators and the area and stratigraphic unit under study (see reviews by Dodson et al., 1980, and Demko and Parrish, 1998). Previous controversy over the paleoclimate of the Morrison Formation largely resulted from conflicting interpretations of the associated flora. Reevaluation of plant taphonomy and taxonomy led Parrish et al. (2004) to conclude that the paleoclimate was warm, seasonal, and semiarid, with the climate changing slightly from dry semiarid throughout most of Morrison deposition to humid-semiarid near the end of Morrison deposition. The semiarid climate in Morrison time probably exerted a primary

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Figure 3. Upper Paleozoic through Upper Jurassic stratigraphy of the Colorado Plateau region. Regional unconformities from Pipiringos and O’Sullivan (1978).

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A Rhizoliths--small and large cf. Planolites isp. Horizontal striated burrows Quasi-vertical striated burrows Adhesive meniscate burrows Ant nests--all nest types Camborygma eumekenomos

cf. Cylindrichum isp. Horizontal U-tubes Vertical Y-tubes cf. Planolites isp. Camborygma litonomos Oxbow lakes/ponds C. airioklados Bivalve traces Channel Gastropod traces Sauropod tracks Ornithopod tracks Theropod tracks

cf. Celliforma--all nest types cf. Rosellichnus isp. Coprinesphaera isp. Various cavities in wood Borings in bones Teethmarks in bones Termite nests--all types

Rhizoliths--small Fuersichnus isp. cf. Planolites isp.

Distal Floodplain Proximal Floodplain Adhesive meniscate burrows Camborygma eumekenomos cf. Ancorichnus isp. Ant nests--all nest types cf. Planolites isp. Fuersichnus isp. Cocoons Vertical burrows cf. Rosellichnus isp. Coprinesphaera isp. cf. Celliforma--all nest types Paleobuprestis isp. Paleoscolytus isp. Various cavities in wood Borings in bones Teethmarks in bones Tektonargus kollospilas Rhizoliths--small and large Tree trunk steinkerns Termite nests--all types

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Levee/Bank Rhizoliths--small and large Adhesive meniscate burrows cf. Ancorichnus isp. Camborygma litonomos C. airioklados cf. Celliforma--solitary Steinichnus isp. Steinichnus isp.--branched Horizontal striated burrows Quasi-vertical striated burrows Termite nests--rhizolith specific

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Figure 4. Trace fossils assemblages characteristic of proximal-to-distal alluvial deposits in typical Mesozoic continental settings. (A) Generalized spatial distribution of continental trace fossils in alluvial settings. (B) Trace fossil assemblages in alluvial deposits in relationship to sedimentation rate and hydrology in channel and overbank settings.

Mesozoic lakes of the Colorado Plateau control on plant recruitment in the environment, which may have resembled a modern savannah. Ground cover in the floodplain regions was predominantly herbaceous, whereas woody vegetation was limited largely to riparian environments (Parrish et al., 2004). The distribution of rhizoliths, with larger ones limited to riparian environments and smaller ones locally present across the floodplain, is consistent with the plant taphonomic inferences drawn by Parrish et al. (2004). The tiering of above- and below-ground trace-making organisms in Morrison deposits, in conjunction with other paleontologic, sedimentologic, stratigraphic, isotopic, and paleogeographic data, suggests spatial and temporal variations associated with tropical wet-dry to Mediterranean-type climates from the southern to the northern part of the Morrison basin (Hasiotis and Demko, 1996; Hasiotis, 2004). Traces of crayfish, termites, ants, bees, beetles, soil bugs, and plants are the preserved products of the water balance in paleosols that record the relation between annual precipitation inputs, solar radiation, evapotranspiration losses, and soil moisture changes during the Late Jurassic. For example, the limited depth, restricted distribution, and low abundance of crayfish burrows in the Salt Wash and Recapture Members suggest that rivers were effluent seasonally and fed the local saturated zone. Burrows with depths of 1–2 m are present close to paleochannels and in very proximal extra-channel environments that were weakly modified by pedogenesis. Termite nests, from <1 to >30 m in depth, indicate shallow to deep saturated zones in proximal to distal alluvial and eolian-derived deposits in the Salt Wash, Recapture, and Brushy Basin Members. Most nests occur in the shallow subsurface in well drained and oxygenated substrates (unsaturated zone) weakly modified by pedogenesis, with fewer and fewer galleries and chambers (fungal gardens, storage, and waste disposal) found deeper in the paleosols. Seasonality of Morrison climate is further substantiated by analysis of annual growth bands in freshwater unionid bivalves from lacustrine beds and fluvial deposits (Good, 2004). Bivalve faunal associations suggest that some of the streams were perennial, although most of the streams in the depositional basin were intermittent. Extended periods of drought might account for some of the famous dinosaur death assemblages (e.g., Carnegie Quarry at Dinosaur National Monument); however, the ecosystem simultaneously sustained some of the most unusual life forms that ever roamed the planet. The semiarid climate interpreted for the Colorado Plateau raises issues about the availability of food and water resources for the large herbivorous dinosaurs (see Engelmann et al., 2004). To survive in such a resource-limited landscape, dinosaurs probably developed various adaptations. Engelmann et al. (2004) discuss how large size conferred an adaptive advantage to sauropods for traveling long distances, enabling the necessary migrations to resources across the landscape with a seasonally dry, semiarid climate. Similarly, scaling effects of large body size make large herbivores efficient relative to their size because they need proportionately less food, and food of lesser quality, than smaller herbivores (Engelmann et al., 2004). In sauropods, differences in dentition and range of neck movement have been used to

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infer styles of resource partitioning, a strategy suitable for survival in a resource-limited environment (Engelmann et al., 2004). A viable alternative interpretation of available evidence suggests a tropical wet-dry climate in the southern part of the Morrison depositional basin and a transition to a Mediterranean-type climate in the northern part of the basin close to or at the edge of the Late Jurassic seaway (Hasiotis, 2004). Throughout Morrison time these paleoclimates fluctuated between seasonally drier and wetter years, similar to climates today (Lydolph, 1985). This fluctuation probably included extreme years with either extended periods of drought (P/E < 1) or precipitation (P/E > 1). The members of the Morrison Formation record high spatial heterogeneity produced by a mosaic of hydroclimates coupled with environments that included dune fields in transitional marine, alluvial, and lacustrine landscapes (Windy Hill, Tidwell, Bluff, and Recapture Members), rapidly aggraded to topographically dissected mixed alluvial landscapes (Salt Wash, Westwater Canyon, Brushy Basin Members), and freshwater to alkaline palustrinelacustrine systems (Tidwell and Brushy Basin Members). These Late Jurassic settings are analogous to modern climates that dominate the African savanna from ~14°N to 5°N latitude and 2°S to 22°S latitude. Modern environments in tropical wet-dry climates contain herds of megaherbivores (elephants, rhinoceros, wildebeests, zebra, and gazelle), predator-scavengers (several types of cats, hyenas, and wild dogs), groups of smaller vertebrates (various birds and rodents), perennial freshwater organisms (fish, crabs, clams, and snails), vast numbers of insects with varying degrees of eusocialism, as well as a variety of plants. This biodiversity is shaped by climate and supports the total biomass through a nutrient and energy cycle robust enough to maintain the ecosystem (e.g., Odum, 1971; Aber and Melillo, 1991; Martinez et al., 1999). Assuming that Jurassic plants, invertebrates, and vertebrate trophic groups such as herbivores, megaherbivores, and predators had water, temperature, and nutrient requirements with ranges of feeding behaviors proportional to their size and physiology comparable to those of extant African biota, then the diversity and distribution of Jurassic continental biota was likely similar to analogous environmental and ecologic settings in modern tropical wet-dry climates. Therefore, the ichnofossils, body fossils, sedimentary facies, paleosols, and geochemical and isotopic signatures of the deposits indicate spatial drier to wetter environmental, hydrologic, and climatic settings within a relatively short distance across the Morrison landscape from any one position at any given time (Hasiotis, 2004). Mesozoic Lake Basin Types Facies associations within members of the Upper Triassic Chinle Formation and Upper Jurassic Morrison Formation exhibit many of the characteristic stacking patterns of lake basin fill under different paleoclimatic settings. Lithofacies representative of fluvial-lacustrine, fluctuating profundal, and evaporative associations are recognized generally on objective physical, chemical and biological criteria in the Monitor Butte Member of

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the Chinle Formation, as well as in the Tidwell and Brushy Basin Members of the Morrison Formation. Hydroclimatic controls during the Late Triassic were dominated by a tropical monsoonal climate that produced varying amounts of sediment through time. Hydroclimatic controls during the Late Jurassic were dominated by a tropical wet-dry climate that experienced periods of seasonally greater and lesser amounts of precipitation through time. Continental trace fossils provide valuable insights into details of deposits interpreted as lake margin, lake plain, palustrine, fluvial, and floodplain paleoenvironments that are often excluded in conventional paleohydroclimate analyses (Hasiotis, 1998, 2004). Trace fossil analysis is especially useful when combined with broader-scale observations of stratal geometries, geochemistry, lithology, and pedogenic features in a lake-basin–type framework. Better integrated analyses of continental alluvial, palustrine, and lacustrine deposits provide more accurate paleoenvironmental reconstructions, including that of the paleohydrologic and paleoclimatic settings during deposition and pedogenic modification of sediments in the early and middle Mesozoic. Please note: any persons wishing to conduct geologic field trips or investigations on the Navajo Reservation, including visiting some of the stops described in this guide, must first apply for and receive a permit from the Navajo Nation Minerals Department, P.O. Box 1910, Window Rock, Arizona 86515, USA. No samples of any kind may be collected while in the National Parks, Recreation Areas, or Navajo Nation Reservation lands without written permission. This includes rock, sediment, vegetation, and cultural materials. Any and all archaeological materials encountered must be left alone and left behind. You may take pictures and leave behind footprints. IMPORTANCE AND GENERAL OVERVIEW OF FIELD TRIP The Upper Triassic Chinle Formation and the Upper Jurassic Morrison Formation are unique in the Mesozoic succession of the Colorado Plateau in that they contain both rich records of continental paleoecosystems and abundant, diverse sedimentary paleoclimatic and paleoenvironmental indicators. Through interpretations of these sedimentary successions and paleontological assemblages, we know that a mosaic of hydrological, ecological, and edaphic conditions characterized both the ancient Chinle and Morrison landscapes. The record of these soils, streams, and lakes of Mesozoic tropical western Pangaea and North America give a glimpse into a dramatic period of Earth history, which included the continued recovery of the continental ecosystem from the Permian-Triassic mass extinction, the appearance and ascendancy of the dinosaurs, and the breakup of a supercontinent. Further evaluation of fluvial-lacustrine deposits in the Colorado Plateau region has enormous potential for further refining our understanding of Mesozoic paleoenvironments and hydroclimatic evolution of the Rocky Mountain orogen and its associated sedimentary basins. This field trip focuses on the lacustrine paleoenvironments within these landscape mosaics and the details of their sedimentary

facies, stratal architecture, trace and body fossil assemblages, and regional paleogeography. However, these archives comprise only part of the rich and fascinating record of this scenic wonderland within the Colorado Plateau. This trip may be initiated from Green River or Moab, Utah. The itinerary outlined here commences from Green River, Utah, and ends at Chimney Rock, located west of Capitol Reef National Park near Torrey, Utah. The general trip route is summarized in Figure 1, and GPS coordinates are provided for some specific landmarks useful for navigation and outcrops (UTM and latitude/longitude in degree decimal on the datum WGS 1984). Many of the stops are accessed via unpaved and unsigned roads. Access to the Day 2 stops, in particular, requires a high clearance and/or four-wheel drive vehicle and suitable weather conditions. As always, good judgment must be used in evaluating local road conditions. FIELD TRIP STOPS Day 1: Saline-Alkaline Lake, Wetland, and Sandy, Ephemeral Fluvial Channel Lake-Margin Deposits of the Brushy Basin Member of the Upper Jurassic Morrison Formation near Moab, Utah, and Four Corners Area Introduction Total mileage for Day 1: ~242 mi (389.5 km). Field stops for Day 1 focus on the upper Brushy Basin Member of the Morrison Formation, interpreted as an alkaline-saline evaporative wetland-lake deposit named Lake T’oo’dichi’ (Navajo for ‘‘bitter water”) (Turner-Peterson, 1987; Turner and Fishman, 1991). Lake T’oo’dichi’ may be the largest and oldest alkaline-saline wetland-lake system described from the geologic record (TurnerPeterson, 1987; Turner and Fishman, 1991). During the Late Jurassic, the ancestral Uncompahgre Uplift imposed a barrier to rivers and shallow, eastward-flowing groundwater that discharged into the San Juan–Paradox Basin on the upstream side of the uplift (Fig. 5). This closed hydrologic setting was necessary for development of a sizeable palustrine-lacustrine environment that persisted for ~2 m.y., based on 40Ar/39Ar dates on minerals from altered ash beds intercalated with the lacustrine deposits (Kowallis et al., 1998). The inputs of silicic volcanic ash were delivered by prevailing winds from a system of calderas located to the west and southwest of the basin. A distinctive lateral hydrogeochemical gradient, indicating increasing salinity and alkalinity in the pore waters, altered the ash to a variety of authigenic minerals in concentric zones defined within the basin (Fig. 6). The basinward progression of diagenetic mineral zones is smectite → clinoptilolite → analcime ± potassium feldspar → albite (Turner and Fishman, 1991). The groundwater-fed wetlands were shallow and evaporated frequently to dryness. Scarce laminated gray mudstone beds record distinct pulses of freshwater lacustrine deposition that resulted from intermittent streams carrying detritus into the basin (Dunagan and Turner 2004). The zeolites, evaporites, and

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authigenic mineral assemblages of Lake T’oo’dichi’ are similar to such Holocene saline-alkaline wetland and lake environments as Lake Magadi, Kenya, and Teels Marsh, Nevada (Surdam and Sheppard, 1978) and Pleistocene Lake Tecopa, California (Sheppard and Gude, 1968, 1986). Planned field stops examine facies of the Upper Brushy Basin Member across a western transect of Lake T’oo’dichi’ (Figs. 5 and 6). Stop 1.1 is along the northern pinch-out, Stop 1.2 is within the clinoptilolite-dominated basin center, and Stop 1.3 is in the analcime zone of the lake, where it was influenced by flashy, ephemeral stream input. Stop 1.1: Courthouse Draw Route and location: After departure from Green River, take I-70 east to Crescent Junction, Utah, and then U.S.-191 (also signed as UT-163) south toward Moab, Utah. Stop 1.1 is off the left side of the road at ~12S 612702 E 4285394 N (38.71188°, −109.70377°). The Brushy Basin Member of the Morrison Formation is characterized by major lateral facies changes from fluvial sandstone beds to low-energy lacustrine sediment and deposits of playa-lake environments containing thick intervals of volcanic tuffs and variegated silty and sandy mudstone overlain by wellsorted and strongly cemented sandstone beds (Turner-Peterson et al., 1986; Bell, 1986). A clinoptilolite-heulandite zeolite mineral assemblage occurs in these sediments deposited along the playa margin. The zeolites formed below the sediment-water interface in these saline-alkaline environments. Near Moab, Utah, in the region northwest of Arches National Park, the contours of ancient Lake T’oo’dichi’ apparently bend in a northward direction toward the area of the present-day Salt Valley anticline (Fig. 6). This suggests that a topographic depression existed in this same area during deposition of the lacustrine sediments during Morrison time. Stop 1.2: Montezuma Creek Route and location: To reach the stops in the Four Corners region, take U.S.-191/UT-163 south toward Monticello and Blanding, Utah. South of Blanding, take UT-262 (turn off at 12S 634119 E 1413583 N; 37.42936°, −109.48344°) east and south to a dirt road turnoff to the right at 12S 648852 E 4132213 N (37.32468°, −109.31927°). The outcrop stops are ~0.5 mi east on this road at 12S 618157 E 4131900 N (37.3220°, −109.3279°). This stop examines the upper part of the Brushy Basin Member in the Morrison Formation within the clinoptilolite diagenetic mineral zone of ancient Lake T’oo’dichi’. The observable ledgeforming units (2–50 cm thick) are tuffs, comprised of altered silicic volcanic ash, that contain a variety of authigenic minerals including mixed-layer illite-smectite, clinoptilolite, analcime, potassium feldspar, albite, quartz, chalcedony, chlorite, kaolinite, barite, calcite, and dolomite (Turner and Fishman, 1991) (Fig. 7). The “popcorn surface texture” results from the weathering of local smectite clays into bentonite (Fig. 8A) and creates the distinctive rounded slopes. The 105-m section of the Brushy Basin

CD 1-1 MC 1-2

BD 1-3

Figure 5. Paleogeographic map of the western United States during Kimmerigian time (from Turner and Peterson, 2004) showing relative locations of Stop 1.1 (Courthouse Draw [CD]), Stop 1.2 (Montezuma Creek [MC]), and Stop 1.3 (Beclabito Dome [BD]).

Member at Montezuma Creek has yielded Late Jurassic ages (149.4 ± 0.7–145.2 ± 1.2 Ma) based on 40Ar/39Ar ages on plagioclase and sanidine grains (Kowallis et al., 1991, 1998). Locally, a brown fossiliferous interval is marked by a laminated, carbonaceous claystone and mudstone unit of palustrine origin and is interbedded with zeolitic tuffs. The laminated claystone preserves plant fossils, including the leaves of the ginkophyte Czekanowskia (Fig. 8B), ferns, and cycads, trunks and stems of conifers and horsetails, and a low-diversity invertebrate fauna including conchostracans (Ash, 1994; Ash and Tidwell, 1998) and is interpreted to be associated with an interval of freshwater lacustrine deposition. This unit grades upward into a clinoptilolite-bearing unit, signaling a return to alkaline-saline palustrine deposition.

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Extent of Lake T?oo?dichi?

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Figure 6. Paleogeographic map of Lake T’oo’dichi’ (from Dunagan and Turner, 2004) showing locations of Day 1 stops relative to mapped diagenetic mineral zones.

Diagenetic mineral zones in tuffs

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Figure 7. (A) Outcrop photo of Montezuma Creek exposure of the Brushy Basin Member of the Morrison Formation (cl—clinoptiolite zone; an—analcime zone). (B) Measured section (from Dunagan and Turner, 2004) of exposure showing zeolite zones within altered tuffs.

Mesozoic lakes of the Colorado Plateau Stop 1.3: Beclabito Dome, New Mexico Route and location: Departing Montezuma Creek for Stop 1.3, take UT-262 east ~17 mi (27.4 km). The route will be resigned CO-41. Continue to drive another 9.5 mi (15.3 km), and turn left on CO-60, then continue driving another 11.2 mi (18 km). Turn left on U.S.-64/AZ-504 another 6.7 mi (10.8 km) (note: the route becomes NM-504). Continue past the town of Beclabito, New Mexico, for Stop 1.3, ~1.4 mi (2.3 km) ahead, on a dirt road to the right at ~12S 67867 E 4077440 N (36.82597°, −108.99656°). The outcrops are located along the wash ~300 m SW of the parking area. This stop along margins of underfilled lake T’oo’dichi’ examines a crossbedded sandy facies derived from ephemeral streams that prograded into the playa system. To reach the outcrops exposed along the wash, head on foot southwest ~0.3 km. Shiprock, a distinctive, Tertiary-age volcanic diatreme, is visible in the distance to the south. Playa aggradational phases are marked by matrix-supported sheetflood deposits with rip-up clasts of tuffaceous material (Fig. 8C). The outcrop has a characteristic spotted texture due to alteration of analcime, other zeolites, and associated minerals. Another diagnostic lithofacies is characterized by an agglomeration of cm-scale iron-rich pellets (Fig. 8D). Most of the palustrine-lacustrine deposits within the upper Brushy Basin Member do not preserve much macroscopic evidence of bioturbation, suggesting high rates of sedimentation, high alkalinity-salinity, anoxic bottom water conditions, or modification by authigenic mineral processes. Abundant but monospecific burrows are present in some thin beds, indicating the presence of only one type of burrowing organism (Hasiotis, 2004). This low diversity ichnofossil assemblage is consistent with harsh environments. Shallow rhizoliths and bioturbation are present in some of the tuffs, indicating that terrestrial vegetation grew locally as part of incipient soil formation during periods of subaerial exposure (Fig. 8E). This exposure was most likely associated with low water levels when the wetland-lake complex evaporated to dryness. The most striking ichnofossil assemblage at this locality includes subterranean termite nests preserved in sandstone deposited by fluvial processes (Fig. 9). The termite nests indicate that the substrates became unsaturated as the saturated zone moved downward and the fluvial deposits became well drained. To return to Bluff, take U.S.-64/AZ-504/NM-504 west ~7 mi (11.3 km). This road will then become U.S.-160; proceed another 28.2 mi (45.4 km). Turn right onto U.S.-191/UT163 and continue driving ~33 mi (53.1 km) to Bluff, Utah, for overnight lodging. Day 2—Fluvio-lacustrine Deposits in the Upper Triassic Chinle Formation in Blue Notch Canyon and North Wash, Glen Canyon National Recreation Area Note: High-clearance two-wheel drive vehicles or fourwheel drive vehicles AND good weather are required for safe access to the Day 2 sites.

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Figure 8. Features in the upper Brushy Basin Member at Montezuma Creek and Belcabito Dome: (A) “Popcorn” weathering of smectite-rich mudstone. (B) Fossil Czekanowskia leaves. (C) Analcime-rich mudstone rip-up clasts. (D) Iron-rich pellets in sandstone. (E) Rootlets within an immature paleosol. Lens cap in photos is 52 mm in diameter.

Introduction Total mileage for Day 2: ~170 mi (273 km). Day 2 focuses on the fluvio-lacustrine deposits of the Upper Triassic Chinle Formation of Glen Canyon National Recreation Area. The majority of the day will be spent in Blue Notch Canyon, where the focus will be on the lacustrine facies preserved in the Monitor Butte Member. Panoramic views combined with excellent exposure permit the recognition and delineation of high-frequency sequences and parasequences from the fluvio-lacustrine successions. From Blue Notch, the route continues to North Wash to demonstrate the lateral variability of the Late Triassic landscape by correlating the Monitor Butte lacustrine sediments to chronostratigraphically equivalent paleosols. Stop 2.1: Blue Notch Pass Route and location: Morning departure from Bluff, taking UT-163 westbound to UT-261, and then northbound to UT-95. Travel northwest 32 mi (51.5 km) on UT-95. See Figure 10 for a detailed route map to the geologic stops. To access Blue Notch Canyon, look for the junction of UT-95 with an unpaved road (Bureau of Land Management [BLM] road 206A) on the left

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Figure 9. Comparison of ichnofossil termite nest (A) in the Brushy Basin Member at Beclabito Dome to modern termite nest architecture (B).

(12S 563021 E 4180252 N; 37.76742° −110.28444°); follow this road to the crest of the pass where there is a wide parking area (12S 561161 E 4180492 N; 37.76972°, −110.30556°). This brief stop examines exposures of palustrine limestone and calcareous mudstone within the Petrified Forest Member of the Chinle Formation. These facies represent marginal lacustrine deposition associated with underfilled lake basins that overlie the Monitor Butte and Moss Back Members of the Chinle Formation, the main units that will be examined in the remainder of the stops during the day. The great exposures and spectacular view from the pass present an opportunity to spend some time introducing the stratigraphy of the Chinle Formation as well as contrasting these underfilled lake basin carbonate facies with the overfilled lake basin and fluvial clastic deposits beneath them. Stops 2.2a through 2.2c: Blue Notch Canyon Introduction: Stops 2.2a–2.2c are along BLM road 206A between the pass at Blue Notch and Lake Powell to the west. In Blue Notch Canyon, the stops focus on facies deposited in a

fluctuating profundal lacustrine environment. Blue Notch Canyon exposes the Lower Triassic Moenkopi Formation, the Monitor Butte, Moss Back, Petrified Forest, Owl Rock, and Church Rock Members of the Upper Triassic Chinle Formation, the Lower Jurassic Wingate Sandstone, and Lower Jurassic Kayenta Formation. Stops highlight the lacustrine deposits of the Monitor Butte Member and the associated fluvial, paludal, and palustrine deposits and paleosols preserved in the lower part of the Chinle Formation (Fig. 11). Within the lower part of the Chinle Formation, the top of the Shinarump Member is an extensive lacustrine flooding surface as well as, in some places, a sequence boundary that marks a previous surface of landscape degradation locally forming a flooding surface sequence boundary (FSSB). Greenish-gray mudstone and brown ripple cross-laminated sandstone of the Monitor Butte Member overlie the lacustrine flooding surface. The Monitor Butte Member locally contains abundant, well preserved plant fossils (Ash, 1975; Demko, 1995). Strata within the Monitor Butte Member are interpreted as a highstand systems tract, deposited as

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Figure 11. Correlation and sequence stratigraphy of the lower part of the Chinle Formation in the White Canyon and North Wash area around the northern part of Glen Canyon National Recreation Area. FS—flooding surface; FSSB—flooding-surface sequence boundary; MFS—maximum flooding surface; SB—sequence boundary.

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a prograding lacustrine delta complex that was fed by a rapidly aggrading fluvial system. The top of the Monitor Butte Member is a sequence boundary, marked either by an erosional surface at the base of the discontinuous Moss Back Member sandstone (where present), or a well developed paleosol characterized by distinctive red coloration and carbonate nodules. The Moss Back Member is a trough crossbedded sandstone characterized by large-scale lateral accretion beds. It represents a smaller, inset incised valley cut into the underlying Monitor Butte lacustrine valley fill. The tripartite fluvial-lacustrine-fluvial succession in the lower Chinle valley fill is reminiscent of other such documented lacustrine basin fills as those in rift basins of the Newark Supergroup (Lambiase, 1990; Olsen, 1990; Carroll and Bohacs, 1999).

Pipiringos and O’Sullivan, 1978). Lacustrine deltaic deposits of the Monitor Butte Member sharply overlie the paleosol, indicating this surface represents a combined FSSB (Fig. 13). The deltaic deposits at the base of the Monitor Butte Member consist of thinly bedded, coarsening upward, fine-grained sandstones that exhibit a complex cut-and-fill architecture. These deposits grade upward to coarser-grained trough crossbedded lacustrine shoreline facies. Two major stages of progradation can be identified by tracing the sharp transition from these shoreline deposits, and a laterally equivalent minor paleosol exposed elsewhere in the canyon, back to fluvial-deltaic deposits. The upper portion of the Monitor Butte Member is truncated by a surface of erosion at the base of the Moss Back Member (Fig. 14A), the focus of the next stop.

Stop 2.2a: Coals in the Lower Part of the Monitor Butte Member of the Upper Triassic Chinle Formation Route and location: Continue westbound from the pass into Blue Notch Canyon on BLM road 206a. The road levels off after a relatively steep descent, and parking is available near the stop at a wide switchback (12 S, 559938 E, 4180432 N; 37.76925°, −110.31944°). This short stop examines unique organic-rich lacustrine facies preserved in the Monitor Butte Member (Fig. 12). The Monitor Butte Member unconformably overlies the Shinarump Member and is composed predominantly of olive-gray to greenish gray smectitic mudstone and siltstone, but also contains fine-grained, tuffaceous sandstone, laminated carbonaceous mudstone and shale (Stewart et al., 1972). The Monitor Butte Member ranges from >80 m thick to zero where it onlaps the paleovalley margins (Stewart et al., 1972). Distinctive sedimentary features in the Monitor Butte include (1) delta foreset beds (Fig. 13); (2) thin, broad distributary channels; and (3) contorted and slumped strata (Stewart et al., 1972; Dubiel et al., 1993). The Monitor Butte Member represents deposition in fluvial, lacustrine, lacustrine delta, and paludal environments (Stewart et al., 1972; Dubiel, 1987; Demko, 2003). In the White Canyon area of southeastern Utah, the Monitor Butte Member also contains thin limestones and coal seams (Dubiel, 1983, 1989a) (Fig. 12). The thin coal seams indicate two distinct sedimentary environments: (1) allochthonous, detrital peat deposited in lacustrine delta-front environments; and (2) autochthonous peat interbedded with fine-grained siliciclastic deposits associated with hydric paleosols developed on transgressive lacustrine margins. Stop 2.2a provides a look at the first type of these organic-rich facies preserved in a lacustrine delta.

Stop 2.2c: Moss Back Member of the Chinle Formation Incised Valley Route and location: Turn around, and head east on BLM road 206a. Park on a wide spot on the road at 12 S, 553596 E, 4178048 N. The Moss Back Member overlies the Monitor Butte Member, or overlies older rocks where the Monitor Butte or Shinarump Members are not present, in a belt ~80–120 km wide (Blakey and Gubitosa, 1983) from northern New Mexico and southwestern Colorado to southeastern Utah. The Moss Back Member ranges from 50 m thick in the White Canyon area of southern Utah and pinches out along the margins of the paleovalley (Stewart et al., 1972). The lower part of the Moss Back Member is characterized by large-scale trough and planar crossbedded, medium-grained sandstone with interbedded carbonate nodule and extrabasinal

Stop 2.2b: Clinoform Delta Foresets in the Monitor Butte Member of the Upper Triassic Chinle Formation Route and location: Continue driving west on BLM road 206a ~0.3 mi (0.5 km) west-northwest of Stop 2.2c. Pull off at a wide spot on the road near 12 S, 553148 E, 4177523 N. This stop features an excellent exposure of a well-developed paleosol characterized by intense bioturbation, color mottling, and vertic features that mark a regional unconformity (Tr-3 of

Figure 12. Coal bed in the lower part of the Monitor Butte Member of the Chinle Formation in Blue Notch Canyon, Glen Canyon National Recreation Area.

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Blue Notch 1 Glen Canyon National Recreation Area Moss Back Mbr.

SB

Monitor Butte Mbr. FS

FSSB SB

Monitor Butte Mbr.

Moenkopi Fm.

Figure 13. Delta front foresets (highlighted by black lines) in the Monitor Butte Member of the Chinle Formation in Blue Notch Canyon, Glen Canyon National Recreation Area. Inset: unconformity paleosol developed on the Moenkopi Formation underlying the Monitor Butte Member of the Chinle Formation, interpreted as a flooding-surface sequence boundary (FSSB) and unconformity. FS—flooding surface; SB—sequence boundary.

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pebble conglomerate near the base of the unit. The upper part of the Moss Back Member is characterized by small- and largescale trough and planar crossbedded, medium- to fine-grained sandstone (Stewart et al., 1972). The Moss Back Member is interpreted to have been deposited in high- and low-sinuosity stream environments. Blakey and Gubitosa (1984) interpreted the sandstones in this unit as low-sinuosity, braided stream deposits, whereas Dubiel (1987), noting the large-scale lateral accretion elements that dominate the internal architecture of the sandstone, viewed them as having been deposited by high-sinuosity, meandering streams. The Moss Back Member outcrops along the southern wall of Blue Notch Canyon and is >30 m thick in the southwestern portion of the canyon near Stop 2.2c and pinches out to the east (Fig. 15). The base of the Moss Back Member is marked by erosional truncation of the underlying Monitor Butte Member. This surface, interpreted as a sequence boundary based on the truncation of underlying strata and a significant basinward shift in facies, can be traced laterally to a well-developed paleosol formed along the Moss Back interfluve. Paleocurrent measurements taken in the canyon (Fig. 14A) suggest the Moss Back

trunk stream flowed westward nearly parallel to the modern canyon wall. After viewing the geologic stops in Blue Notch Canyon, retrace the route east on BLM road 206a back to the paved road, UT-95. Turn left and drive north 20 mi (32.2 km) on UT-95 to Stop 2.3a. Stops 2.3a through 2.3b: North Wash Introduction: These stops demonstrate the stratal architecture and facies relationships within the paleovalley fill. Compared to Blue Notch, the section at North Wash is markedly thinner, and lacustrine facies are absent. Despite the differences between these two sections, spatio-temporal correlations can be made on the basis of paleohydrologic soil indicators. These correlations provide some basis to constrain the facies relationships within the paleovalley (Beer, 2005) (Fig. 11). Stop 2.3a: Depositional Facies Preserved on Paleovalley Margins Route and location: Drive 20 mi (32.2 km) north of the Blue Notch canyon turnoff on UT-95. Take a left into a gravel parking

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Figure 14. Measured section of the lower part of the Chinle Formation. (A) Blue Notch Canyon, Glen Canyon National Recreation area at Stop 2.2b; note paleocurrent rose diagram for Moss Back Member sandstone. (B) Stop-2.3a, North Wash, Glen Canyon National Recreation Area. (C) Chimney Rock in Capitol Reef National Park at Stop 3.4b.

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lot. To reach the outcrop, hike along the road south (back toward Hite) and up the first major side canyon on the east side of the highway (with luck, you can walk directly up the wash starting at the culvert that passes under the highway). The outcrops are at 12 S, 548704 E 4195924 N (37.909555°, −110.445944°). A 20 min hike up the side canyon ends at the outcrop in Figure 16, where a major erosional unconformity marks the base of the Chinle Formation. Access to a nearly complete exposure of the Monitor Butte and Moss Back Members provides a look at high-sinuosity fluvial, palustrine, and paludal deposits, along with paleosols associated with both landscape aggradation and degradation (Fig. 14B). Stop 2.3b: North Wash—High-Sinuosity Lateral Accretion Deposits Route and location: Continue north on UT-95 2.3 mi (3.7 km) to a safe place to pull off the road near the entrance sign for Glen Canyon National Recreation Area. The outcrop is a short distance off the west side of the highway. High-sinuosity fluvial deposits and associated paleosols comprising the top of the Monitor Butte Member are apparent below the Moss Back Member (Fig. 17). These deposits are similar to those found elsewhere in the basin marking the final basinward progradation of the Monitor Butte fluvio-lacustrine system. These depositional and pedogenic facies can be correlated 50 mi (80.5 km) northward to the closest exposures of the Chinle Formation in the San Rafael Swell. Chinle strata below these high-sinuosity deposits onlap the margin of the master paleovalley over this distance, and the lithostratigraphic nomenclature of these units changes from Monitor Butte Member in the Glen Canyon area to the Temple Mountain Member in the San Rafael Swell area. This ends the stops for Day 2. After Stop 2.3, continue north on UT-95. Go left on UT-276, westbound to Ticaboo, Utah, for overnight lodging. Day 3—Balanced-Fill Lake Deposits in the Tidwell Member of the Upper Jurassic Morrison Formation and High-Frequency Sequence Stratigraphy of FluvioLacustrine Deposits in the Lower Part of the Upper Triassic Chinle Formation Introduction Total mileage for Day 3: ~117 mi (188 km). Stops on the morning of Day 3 will examine outcrops of the Tidwell Member of the Upper Jurassic Morrison Formation interpreted as fluctuating-profundal and palustrine deposits in balanced-fill lakes in the Henry Mountains, Waterpocket Fold, and Capitol Reef National Park areas (Fig. 18). Stops in the afternoon of Day 3 will demonstrate the correlation of parasequences, parasequence sets, and high-frequency sequences within overfilled lakes and associated fluvial deposits the lower part of the Upper Triassic Chinle Formation (Shinarump, Monitor Butte, and Moss Back Members) in the Capitol Reef National Park area (Fig. 18).

Stop 3.1: Tidwell Member of the Morrison Formation, Photo Panorama Interpretation Route and location: After a morning departure from lodging at Ticaboo, Utah, drive north on UT-276 7.5 mi (12.1 km) to a pullout on the east side of the road at 12S 530703 E 4179594 N (37.76314°, −110.65142°). This stop will introduce the stratigraphy of the Middle and Upper Jurassic strata of the Henry Mountains, Waterpocket Fold, and Capitol Reef areas, including recognition of the boundaries between the Middle Jurassic Summerville and Upper Jurassic Morrison Formations, and the stratal architecture of the basal Tidwell and overlying Salt Wash Members of the Morrison Formation (Fig. 19). Stop 3.2: Tidwell Member of the Morrison Formation, Measured Section Route and location: Turn around and head back south on UT-276 3.4 mi (5.5 km) and turn right onto an unpaved road at 12S 527593 E 4174751 N (37.71961°, −110.68692°). Continue driving ~5.7 mi (9.2 km), past the uranium mill facilities (but not onto their access road), into Shootaring Canyon to Stop 3.2, near the Tony M Mine (12S 526105 E 4175302 N; 37.72461°, −110.70378°). Shootaring Canyon is also spelled Shootering, and on some older maps it is also known as Shitamaring Canyon. This area was intensely prospected and mined for uranium ore bodies hosted in sandstones in the Salt Wash Member of the Morrison Formation. Outcrops in Shootaring Canyon preserve vertebrate tracks, trackways, and burrows in the Salt Wash Member. In the canyon, floodplain paleosols interpreted as stacked argillisols and entisols, are intercalated with crevasse-splay deposits, levee sandstones, and channel and bar macroforms. This stop is located at a well-exposed interval of the upper part of the Middle Jurassic Summerville Formation and the lower part of the Upper Jurassic Morrison Formation (Fig. 20A), near the recently active (ca. 2001) Tony M uranium mine. Here, the basal Tidwell Member overlies the J-5 regional unconformity (Pipiringos and O’Sullivan, 1978) cut into the Summerville Formation and is marked by a sandstone bed containing a lag deposit of chert granules and pebbles. This unit is a regional marker bed informally called “Bed A.” Here in the Henry Mountains area, Bed A is interpreted as a unit that was deposited in a desert environment characterized by ephemeral streams and low-relief, eolian sand sheets (zibars). A progradational, coarsening-upward succession of mudstone, siltstone, and silty limestone immediately overlies Bed A (Fig. 20A), marking a lacustrine flooding surface and subsequent deposition of fluctuating-profundal lake facies in a balance-filled basin (Fig. 21). Lacustrine trace fossil assemblages preserved in the Tidwell Member vary in composition areally due to the hydrologic conditions, with greatest benthic ichnodiversity in proximal settings (Fig. 22). These results are consistent with benthic diversity data

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Figure 15. Interpreted photo panorama of the edge of the Moss Back paleovalley in Blue Notch Canyon, Glen Canyon National Recreation Area. Note coarsening-upward succession of the deltaic parasequence below the Moss Back unconformity.

North Wash 2 Glen Canyon National Recreation Area

Monitor Butte Mbr.

Moenkopi Fm. Corey for Scale

Figure 16. Interpreted photo panorama of Monitor Butte lacustrine facies filling the paleovalley cut into the underlying Moenkopi Formation in North Wash, Glen Canyon National Recreation Area. Note person for scale.

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Figure 17. Interpreted photo panorama of high-sinuosity fluvial facies at the top of the Monitor Butte Member of the Chinle Formation at Stop 2.3b. Note lateral accretion surfaces dipping to the right.

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Figure 19. Interpreted photo panorama of the upper part of the Middle Jurassic Summerville Formation (Js) and the Tidwell Member (Jmt) and lower part of the Salt Wash Member (Jms) of the Upper Jurassic Morrison Formation at Stop 3.1.

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Figure 20. Measured sections of the upper part of the Middle Jurassic Summerville Formation and the Tidwell Member and lower part of the Salt Wash Member of the Upper Jurassic Morrison Formation. (A) Stop 3.2 in Shootaring Canyon near the Tony M uranium mine. (B) Stop 3.3 off the Notom Road and UT-24 near the Fremont River. See symbol key in Figure 14.

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NR 3-3 SC 3-2

Figure 21. Paleogeography of the western United States during early Kimmerigian time (from Turner and Peterson, 2004) showing relative locations of Stop 3.2 (Shootaring Canyon [SC]) and Stop 3.3 (Notom Road [NR]) in the Tidwell Member of the Morrison Formation.

from modern lakes (Ward, 1992) such that the profundal zone contains mostly ichnofossils interpreted as those constructed by tubificid annelids and chironomid larvae. In general, tiering in lacustrine environments is quite shallow and can be subtle, dominated by hydrophilic traces no greater than 20 cm in depth below the sediment-water interface, with the majority within 5 cm of the paleolake bottom (Fig. 22).

Morrison lacustrine trace fossils do not fit into the Mermia ichnofacies as defined by Buatois and Mángano (1995) and Buatois et al. (1998) to characterize bioturbation in lake deposits. Many of the Morrison trace fossils indicate firm substrates in shallow water with intermittent subaerial exposure. Environments in deeper water settings do not show any of the diversity expected for the purported Mermia ichnofacies; only Planolites and simple ghost U-tubes are present in Morrison sublittoral deposits. Many characteristic ichnotaxa of the Mermia ichnofacies are absent from Morrison lacustrine deposits, with the exception of Cochlichnus and cf. Planolites, which are found in littoral environments. Such simple horizontal, vertical, and U-shaped feeding and burrowing traces as Planolites, Palaeophycus, Arenicolites, and Skolithos are likely to be present in lacustrine ichnocoenoses because similar forms can be made in any environmental conditions. The paleoenvironmental and paleoecological significance of such structures, however, would be different due to the different conditions found in lacustrine deposited strata. The thick sandstone of the overlying Salt Wash Member sits on a sequence boundary, marked by regional truncation of strata, a significant basinward shift in facies, and an increase in grain size. If time permits, an optional stop will be made further up the main part of the canyon to examine dinosaur footprints and trackways, along with other ichnofossils, in floodplain deposits of the overlying Salt Wash Member of the Morrison Formation (Fig. 23). Stop 3.3: Tidwell Member of the Morrison Formation along the Fremont River Route and location: To proceed toward Stop 3.3, take UT276 west ~11.7 mi (18.2 km) to UT-262 (Burr Trail). Continue on UT-262 north for another 55 mi (88.5 km). At Notom, take the left fork (Notom Road) toward UT-24. It is 3.4 mi (5.5 km) from the junction of Burr Trail (UT-262) and Notom Road to UT-24. Stop 3.3 is at outcrop directly east of the intersection of Notom Road and UT-24, on the south side of UT-24 at 12S 498923 E 4190134 N (37.85867°, −111.01235°). Here the Tidwell Member of the Morrison Formation is dominated by marginal lacustrine and palustrine deposits, including rooted and trampled (dinoturbated) calcareous mudstone and calcrete layers and nodules (Fig. 20B). At this locality “Bed A” is much coarser than at the previous stop in Shootering Canyon and is characterized by massive to crude bedding, poor sorting, and a matrix-supported fabric and texture. “Bed A” here is interpreted to have been deposited in an ephemeral stream environment characterized by significant debris flow deposition and eolian modification. Stops 3.4a through 3.4b: Capitol Reef National Park Introduction: Traveling west from Stop 3.2, state highway UT-24 cuts through the Waterpocket Fold into the Mesozoic section. Here, all three sequence-bounding unconformities in the lower portion of the Chinle (Shinarump, Monitor Butte, Temple Mountain, and Moss Back Members), which define three periods

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Horizontal striated burrows Rhizoliths--small and large cf. Celliforma--all nest types Vertebrate tracks cf. Planolites isp. Various cavities in wood Termite nests--rhizolith associated Ant nests--all nest types Termite nests--all types cf. Celliforma--all nest types Alluvial-Supralittoral Lacustrine Rhizoliths--small and large Deep Water Table Settings cf. Planolites isp. Ant nests--all types X Supralittoral Lacustrine Dune fields

cf. Cylindrichum isp. Horizontal U-tubes Vertical Y-tubes cf. Planolites isp. Camborygma litonomos C. airioklados Bivalve traces Channel Gastropod traces Sauropod tracks Ornithopod tracks Theropod tracks

Littoral Lacustrine Rhizoltihs--small Adhesive meniscate burrows Camborygma litonomos cf. Cylindrichum isp. C. airioklados cf. Planolites isp. Steinichnus traces Gastropod traces X' Stromatolites Stromatolites with borings Theropod tracks cf. Celliforma--solitary Shallow U-tubes Fuersichnus isp. Pteraichnus tracks Horizontal U-tubes Cochlichnus isp. Sauropod tracks Ornithopod tracks Kouphichnium traces

Alluvial-Supralittoral Lacustrine Shallow Water Table Settings Rhizoliths--small and large Horizontal striated burrows Quasi-vertical striated burrows Adhesive meniscate burrows Ant nests--all nest types Camborygma eumekenomos cf. Ancorichnus isp. Termite nests--all types

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Figure 22. Trace fossils assemblages characteristic of supralittoral to profundal lacustrine deposits in typical Mesozoic continental settings. (A) Generalized spatial distribution of continental trace fossils in lacustrine settings. (B) Onshore-offshore transect of supralittoral to profundal lacustrine systems showing environmental gradients. C. Trace fossil assemblages characteristic of lacustrine subenvironments. At—ant nest; An—Anchorichnus; Ca—Cambroygma; Ce—Celliforma; F—Fuersichnus; G—gastropod trail; Hu—horizontal u-tubes; Km—Kouphichnium; P—Planolites; Pts—pterosaur scratch marks; Rh—rhizoliths; Rl—Rosellichnus; Sa—sauropods tracks; T/Rh—termite/rhizolith nest; Tm—termite nest; Ut—ushaped tubes; Uts—shallow u-shaped tubes; Vb—vertebrate burrows. Trace fossil illustrations and box diagrams are schematic and not to scale.

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Figure 23. Sauropod footprints (marked by arrows) in alluvial facies of the Salt Wash Member of the Upper Jurassic Morrison Formation, Shootaring Canyon.

of incision and subsequent valley fill, can be seen (Fig. 24). The initial period of landscape degradation is marked by interfluve paleosols and truncation of the underlying Moenkopi Formation, creating a paleovalley that constrains the deposition of these four members, and defining the basal sequence-bounding unconformity (Demko, 2003; Beer, 2005). The first paleovalley fill, represented by the Shinarump Member, is interpreted as a confined, sandy, low-sinuosity river system. A second period of incision is marked by truncation of the Shinarump and correlative paleosols and pedogenically modified strata (Demko, 2003; Beer, 2005). These features define the second sequence-bounding unconformity. Above this unconformity, mudstones and sandstones of the Monitor Butte and correlative Temple Mountain Members were deposited, representing fluvio-lacustrine deposition and vertisol development in a high-sinuosity river system (Beer, 2005). Truncation of the Monitor Butte and Temple Mountain Members and the Moenkopi Formation, along with interfluve pedogenesis, mark the final cut-and-fill sequence occupying the master paleovalley. The high- and low-sinuosity fluvial deposits of the Moss Back Member overlie the surface of landscape degradation. The first complete section of Chinle Formation is exposed across from “Guy Smith’s Place” in historic Fruita, where a white-weathering paleosol developed at the top of the Lower Triassic Moenkopi Formation marks the base of the Monitor Butte Member. Compared to the deposits in Blue Notch Canyon, the lower part of the Chinle Formation (i.e., the Shinarump, Monitor Butte, and Moss Back Members) is significantly thicker in Capitol Reef National Park. These sediments were deposited closer to the center of the master paleovalley, and as a result a significant thickness of Shinarump basal valley-fill has been preserved (Fig. 25). Stop 3.4a: Panorama Point—Capitol Reef National Park Route and location: After Stop 3.3, head west 8.7 mi (14 km) on UT-24 toward Torrey, Utah. See Figure 18 for a detailed map of the next geologic stops. Turn left off UT-24

and follow the signs for Panorama Point at signed pulloff; 12S 0474148 E 4239944 N (38.30722°, −111.29569°). From the vantage point at the Goosenecks overlook 1.7 mi west of the Capitol Reef National Park Visitor Center, we clearly see the nature of the surfaces bounding the Shinarump Member, as well as the stratal architecture of Monitor Butte Member deltaic clinoforms. The Shinarump Member is bounded above and below by regionally extensive unconformities. These surfaces are often marked by well-developed paleosols or by erosional truncation of underlying strata. Erosional truncation of the underlying Moenkopi Formation marking the Tr-3 unconformity can be seen at outcrop scale in many places on the Colorado Plateau; however, truncation of the upper part of the Shinarump Member of the Chinle Formation is often much more subtle. At the Goosenecks Overlook, several meters of truncation is clearly visible and the surface is overlain by Monitor Butte Member fluvio-lacustrine deposits (Fig. 25). This surface is interpreted as a FSSB. Stop 3.4b: Chimney Rock—Capitol Reef National Park Route and location: To reach Stop 3.4b, return to UT-24 and continue west 0.5 mi (0.8 km). Turn right following a sign for the Chimney Rock Trailhead. At 0.5 mi (0.8 km) west of Panorama Point, turn right into the parking lot just north of the highway. A short hike heading northwest from the parking lot at Chimney Rock provides access to the Shinarump, Monitor Butte, and Moss Back Members of the Chinle Formation (Figs. 14C, 25, and 26). The lower two-thirds of the Monitor Butte Member consist of two distinct phases of lacustrine progradation separated by a distinctive flooding surface. The overlying portion of the Monitor Butte is characterized by paleosols associated with landscape aggradation, pedogenically modified crevasse splay deposits, and lateral accretion sets. These depositional facies represent a highsinuosity fluvial system that prograded into the lake. The contact between the Monitor Butte and Moss Back Members is marked by a well-developed paleosol, which forms a resistant red-orange cliff in the outcrop. This stop ends the geologic field guide. To continue to Salt Lake City, Utah, continue west and then north on UT-24 to Salina, Utah, and then take either U.S.-50 or UT-28 to I-15N. It is ~220 mi (354 km) to Salt Lake City. ACKNOWLEDGMENTS We thank the Navajo Nation and the National Park Service for their support of our continuing research. TMD acknowledges financial support from the University of Minnesota–Duluth, the National Science Foundation, and the National Park Service. KN acknowledges financial support from the Royal Society, the National Aeronautics and Space Administration Global Change/ Mission to Planet Earth, and Brasenose and St. Catherine’s Colleges, University of Oxford. JJB thanks Corey Wendland and Ryan Erickson for help with field work and acknowledges financial support provided by the Colorado Scientific Society. STH thanks the American Association of Petroleum Geologists, the

Mesozoic lakes of the Colorado Plateau

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Figure 24. Correlation and sequence stratigraphy of the lower part of the Chinle Formation in and around the central part of Capitol Reef National Park. FS—flooding surface; FSSB—flooding-surface sequence boundary; SB—sequence boundary.

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Figure 25. Interpreted photo panorama of the lower part of the Chinle Formation from Panorama Point in Capitol Reef National Park (Stop 3.4a). Note incision of the top of the Shinarump Member and clinoforms of delta foresets in the Monitor Butte Member.

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Figure 26. Interpreted photo panorama of the lower part of the Chinle Formation at Chimney Rock in Capitol Reef National Park (Stop 3.4b). MFS—maximum flood surface.

Geological Society of America, and the Petrified Forest Museum Association for funding. Collectively, we thank Fred Peterson and Christine Turner for all of their help and support, and our introduction to the mysteries of Lake T’oo’dichi’. REFERENCES CITED Aber, J.D., and Melillo, J.M., 1991, Terrestrial ecosystems: Philadelphia, Saunders College Publishing, 429 p. Ash, S.R., 1967, The Chinle (Upper Triassic) megaflora of the Zuni Mountains New Mexico: New Mexico Geological Society Guidebook, 18th Field Conference, p. 125–131. Ash, S.R., 1972, Plant megafossils of the Chinle Formation, in Breed, W.J., and Breed, C.S., eds., Investigations in the Triassic Chinle Formation: Museum of Northern Arizona Bulletin, v. 47, p. 23–43. Ash, S.R., 1975, The Chinle (Upper Triassic) flora of southeastern Utah: Four Corners Geological Society Guidebook, 8th Field Conference, Canyonlands, p. 143–147. Ash, S.R., 1980, Upper Triassic floral zones of North America, in Dilcher, D.L., and Taylor, T.N., eds., Biostratigraphy of fossil plants: Stroudsburg, Pennsylvania, Dowden, Hutchinson and Ross, p. 153–170.

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Peterson, F., 1994, Sand dunes, sabkhas, streams, and shallow seas: Jurassic paleogeography in the southern part of the Western Interior basin, in Caputo, M.V., Peterson, J.A., and Franczyk, K.J., eds., Mesozoic systems of the Rocky Mountain region, USA: Denver, Colorado, Rocky Mountain Section, Society of Economic Paleontology and Mineralogy, p. 233–272. Peterson, F., and Turner-Peterson, C.E., 1987, The Morrison Formation of the Colorado Plateau: Recent advances in sedimentology, stratigraphy, and paleotectonics: Hunteria, v. 2, p. 1–18. Pipiringos, G.N., and O’Sullivan, R.B., 1978, Principal unconformities in Triassic and Jurassic rocks, Western Interior United States—a preliminary survey: U.S. Geological Survey Professional Paper 1035-A, 29 p. Riggs, N.R., Lehman, T.M., Gehrels, G.E., and Dickinson, W.R., 1996, Detrital zircon link between headwaters and terminus of the Upper Triassic Chinle-Dockum paleoriver system: Science, v. 273, p. 97–100. Schudack, M.E., Turner, C.E., and Peterson, F., 1998, Biostratigraphy, paleoecology and biogeography of charophytes and ostracodes from the Upper Jurassic Morrison Formation, Western Interior, USA: Modern Geology, v. 22, p. 379–414. Sheppard, R.A., and Gude, A.J., 1968, Distribution and genesis of authigenic silicate minerals in tuffs of Pleistocene Lake Tecopa, Inyo County, California: U.S. Geological Survey Professional Paper 597, 38 p. Sheppard, R.A., and Gude, A.J., 1986, Magadi-type chert—a distinctive diagenetic variety from lacustrine deposits, in Mumpton, F.A., ed., Studies in diagenesis: U.S. Geological Survey Bulletin 1578, p. 335–345. Surdam, R.C., and Sheppard, R.A., 1978, Zeolites in saline, alkaline-lake deposits, in Sand, L.B., and Mumpton, F.A., eds., Natural zeolites—Occurrences, properties, use: Pergamon, New York, Pergamon Press, p. 145–174. Stewart, J.H., Poole, F.G., and Wilson, R.F., 1972, Stratigraphy and origin of the Chinle Formation and related Upper Triassic strata in the Colorado Plateau region: U.S. Geological Survey Professional Paper 690, 336 p. Turner, C.E., and Fishman, N.S., 1991, Jurassic Lake T’oo’dichi’: A large alkaline, saline lake, Morrison Formation, eastern Colorado Plateau: Geological Society of America Bulletin, v. 103, p. 538–558, doi: 10.1130/00167606(1991)103<0538:JLTODA>2.3.CO;2. Turner, C.E., and Peterson, F., 1999, Biostratigraphy of dinosaurs in the Upper Jurassic Morrison Formation of the Western Interior, USA, in Gillette, D.D., ed., Vertebrate paleontology in Utah: Utah Geological Survey Miscellaneous Publication 99-1, p. 77–114. Turner, C.E., and Peterson, F., 2004, Reconstruction of the Upper Jurassic Morrison Formation extinct ecosystem—a synthesis: Sedimentary Geology, v. 167, p. 309–355, doi: 10.1016/j.sedgeo.2004.01.009. Turner-Peterson, C.E., 1987, Sedimentology of the Westwater Canyon and Brushy Basin Members, Upper Jurassic Morrison Formation, Colorado Plateau, and relationship to uranium mineralization [Ph.D. dissertation]: Boulder, University of Colorado, 184 p. Turner-Peterson, C.E., Santos, E.S., and Fishman, N.S., editors, 1986, A basin analysis case study; the Morrison Formation, Grants uranium region, New Mexico: American Association of Petroleum Geologists Studies in Geology, v. 22, 391 p. Vail, P.R., Mitchum, R.M., and Thompson, S., 1977, Seismic stratigraphy and global changes of sea level, part 3: Relative changes of sea level from coastal onlap, in Payton, C.E., ed., Seismic stratigraphy—Applications to hydrocarbon exploration: American Association of Petroleum Geologists Memoir 26, p. 63–81. Valdes, P.J., and Sellwood, B.W., 1992, A palaeoclimate model for the Kimmeridgian: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 95, p. 47–72, doi: 10.1016/0031-0182(92)90165-2. Ward, J.V., 1992, Aquatic insect ecology I: Biology and habitat: New York, Wiley, 456 p. Ziegler, A.M., Scotese, C.R., and Barren, S.F., 1983, Mesozoic and Cenozoic paleogeographic maps, in Brosche, P., and Sundermann, J., eds., Tidal friction and the Earth’s rotation, II: Berlin, Springer-Verlag, p. 240–252. Ziegler, A.M., Parrish, J.M., Yao, J.P., Gyllenhaal, E.D., Rowley, D.B., Parrish, J.T., Nie, S.Y., Bekker, A., and Hulver, M.L., 1993, Early Mesozoic phytogeography and climate: Philosophical Transactions of the Royal Society of London, v. 341, p. 297–305.

Printed in the USA

Geological Society of America Field Guide 6 2005

Birth of the lower Colorado River—Stratigraphic and geomorphic evidence for its inception near the conjunction of Nevada, Arizona, and California P. Kyle House Nevada Bureau of Mines and Geology, University of Nevada, Reno, Nevada 89557, USA Philip A. Pearthree Arizona Geological Survey, 416 W. Congress #100, Tucson, Arizona, 85701, USA Keith A. Howard U.S. Geological Survey, Menlo Park, California 94025, USA John W. Bell Nevada Bureau of Mines and Geology, University of Nevada, Reno, Nevada 89557, USA Michael E. Perkins Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112, USA James E. Faulds Nevada Bureau of Mines and Geology, University of Nevada, Reno, Nevada 89557, USA Amy L. Brock Department of Geoscience, University of Nevada, Las Vegas, Nevada 89154, USA

ABSTRACT A detailed record of the late Cenozoic history of the lower Colorado River can be inferred from alluvial and (likely) lacustrine stratigraphy exposed in dissected alluvial basins below the mouth of the Grand Canyon. Numerous sites in Mohave, Cottonwood, and Detrital valleys contain stratigraphic records that directly bear on the mode, timing, and consequences of the river’s inception and integration in the latest Miocene–early Pliocene and its subsequent evolution through the Pleistocene. This field trip guide describes and illustrates many of these key stratigraphic relationships and, in particular, highlights evidence that supports the hypothesis of cascading lake-overflow as the principal formative mechanism of the river’s course downstream from the Grand Canyon. Keywords: Colorado River, Bouse Formation, Chemehuevi Formation, flood, stratigraphy. House, P.K., Pearthree, P.A., Howard, K.A., Bell, J.W., Perkins, M.E., Faulds, J.E., and Brock, A.L., 2005, Birth of the lower Colorado River—Stratigraphic and geomorphic evidence for its inception near the conjunction of Nevada, Arizona, and California, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 357–387, doi: 10.1130/2005.fld006(17). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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INTRODUCTION Many details about the mode, timing, and consequences of the inception of the lower Colorado River have eluded geologists for the past 150 years. Much of the research into the problem has focused on the Grand Canyon because of its spectacular setting, but the dominance of erosion there limits the amount and kind of information that can be gleaned about the river’s history. Valleys downstream from the Grand Canyon, however, contain extensive exposures of a corresponding record of river and tributary deposition. That record is, in turn, superposed on a preriver stratigraphic framework that documents the environments into which the river first flowed. The composite record supports a robust model of lower Colorado River evolution over the past 5–6 m.y. The principal geographic focus of this field trip guide is the Cottonwood Valley–Mohave Valley reach of the lower Colorado River (Fig. 1). Here, the river drains an ~169,300 mi2 (433,000 km2) watershed that has undergone a series of major changes during the late Cenozoic, including: large-scale drainage integration, excavation of numerous canyons upstream, and multiple climatic changes. Aspects of these events are revealed in extensive exposures of late Cenozoic sedimentary deposits and landforms in Cottonwood and Mohave valleys. Recent detailed mapping of these deposits reveals a richer and more complex record than has been previously known (Faulds et al., 2004; House et al., 2002, 2004a; Faulds and House, 2000; Pearthree and House, 2004). Basic Geologic Setting of the Lower Colorado River The lower Colorado River extends from the mouth of the Grand Canyon at the western edge of the Colorado Plateau to the Gulf of California, traversing a series of rugged bedrock canyons separated by relatively broad, elongate alluvial basins typical of the Basin and Range province. The basins are the product of large-scale, regional crustal extension between ca. 25 and 10 Ma (Anderson, 1971; Howard and John, 1987; Spencer and Reynolds, 1989). The river follows a Miocene extensional corridor between Hoover Dam and Parker whose 50–100 km east-west width was doubled by crustal stretching. This extension resulted in tilted fault blocks, many normal faults, and metamorphic core complexes. The modern course of the Colorado River through this extensional terrain is relatively young. Geologic relations at the western end of Grand Canyon indicate that an integrated Colorado River did not exist prior to 5.6–6 Ma (e.g., Lucchitta, 1979; Spencer et al., 2001; Faulds et al., 2002; Howard et al., 2000); but lava flows intercalated with river gravels within ~110 m of its modern level west of Grand Canyon indicate that it was established by 4.7–4.4 Ma (Howard and Bohannon, 2001). Sedimentologic and paleomagnetic data from the river’s early delta in the Salton Trough area indicate its arrival in that area before 4.3 Ma (Johnson et al., 1983). Relations that we have discovered in Cottonwood and Mohave valleys are similar and preclude the presence of a throughgoing major river prior to 5.5 Ma, but require it before 4 Ma.

In the various canyons between Lake Mead and Yuma, deposits of Colorado River gravel occur up to 250 m (800 ft) above the modern river, and discernible bedrock straths and paleochannel features are found at similar and progressively lower levels. Intervening alluvial valleys along the river’s modern course contain similarly high fluvial deposits, flights of younger fluvial terraces, and successions of fluvial scarps. As many as seven prominent terrace levels occur along the river, ranging from the modern floodplain 2–10 m above the active channel to the oldest well-preserved terrace that stands ~110 m above the river in Mohave Valley. Higher and older remnant river deposits provide an important component to understanding the early history of the lower Colorado River. Cottonwood Valley, Pyramid Canyon, and Mohave Valley A detailed geologic record of the river’s late Cenozoic history is preserved along the reach of the lower Colorado River through southern Cottonwood Valley, Pyramid Canyon, and Mohave Valley (Fig. 1). The western edge of this area is bounded by a series of mountain blocks underlain predominantly by Miocene plutonic rocks with a variable cover of Miocene volcanic rocks. The eastern edge is bounded by the Black Mountains, a complex of Proterozoic basement (largely granite and gneiss) with a locally thick cover of Miocene volcanic and minor sedimentary rocks. Mohave and Cottonwood valleys are separated by the “Pyramid hills” (informal name), a relatively low set of rugged bedrock hills that straddle the valley axis between the Black and Newberry mountains. The Pyramid hills are cored by the Davis Dam granite (Faulds et al., 2004), a megacrystic Precambrian granite porphyry. The south end of Mohave Valley abruptly terminates at Topock Gorge, a narrow canyon between the Mohave (east) and Chemehuevi (west) mountains. Mohave Valley is broad and contains extensive exposures of river deposits, late Cenozoic basin fill, and a suite of fluvial landforms. Southern Cottonwood Valley is narrower and is dominated by steeply sloping alluvial fan remnants that flank the shores of Lake Mohave. Post-Miocene tectonic activity in both valleys has been minor. Small faults and some minor tilting have been noted in some late Tertiary basin fill deposits near Laughlin and north of Katherine’s Landing. In southern Mohave Valley, structural perturbation of early to middle Pleistocene alluvial fan deposits is clearly evident on the distal end of the Warm Springs fan complex near Golden Shores, Arizona (Purcell and Miller, 1980; Pearthree et al., 1983). Previous Work on Lower Colorado River Stratigraphy The late Cenozoic alluvial stratigraphy of Mohave and Cottonwood valleys has been studied and described in various levels of detail beginning with the first geologic observations by Newberry in 1857–1858 (in Ives, 1861). Subsequently, a more detailed stratigraphic framework began to evolve from Lee (1908), Blackwelder (1934), Longwell (1936, 1947, 1963), and

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Metzger et al. (1973). The previous work on the river’s alluvial stratigraphic record and underlying basin deposits has provided an essential foundation for our detailed studies in the Laughlin area (Fig. 2). Early reports on the geomorphology of the lower Colorado River (Newberry, 1861, in Ives, 1861; Lee, 1908) emphasized fluvial terraces and described 3–5 distinct levels in Mohave Valley. Lee (1908) developed a model of river evolution involving

three canyon cutting intervals separated by three major aggradation events. He used the term Chemehuevis gravels for the most conspicuous fill deposits along the river, which were associated with his second major aggradation event. He proposed that the aggradational packages were responses to climatic control. Blackwelder (1933) counted 9–12 distinct terrace levels in the Yuma area and tentatively attributed them to climatic changes or eustatic changes in the Gulf of California. Longwell (1936) investigated

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Figure 2. Comparison of various conceptions of the late Cenozoic stratigraphy of the lower Colorado River. Coeval and intervening tributary alluvial fan deposits not shown. Fm—Formation.

the Colorado River deposits in great detail in what is now the Lake Mead area prior to the construction of Hoover Dam. He identified as many as seven levels of alluvial terraces along the river and highlighted the apparent significance of what he renamed the Chemehuevi Formation, which he presumed was a major Pleistocene aggradation package (Fig. 2A). He attributed the river’s stratigraphic record to climate variability, but also proposed a lacustrine model for the Chemehuevi Formation associated with some form of downstream blockage (Longwell, 1947, 1963). In a series of regional hydrogeologic studies, Metzger et al. (1973), Metzger and Loeltz (1973), and Olmstead et al. (1973) evaluated Colorado River deposits from Davis Dam to Yuma and developed the most robust stratigraphic framework for the river’s history up to that time (Fig. 2B). It strongly influenced subsequent studies in the 1970s and was augmented by Lee and Bell (1975), Bell et al. (1978), and Ku et al. (1979), who developed some of the first age-controls on key river deposits and related piedmont geomorphic surfaces using U-series, amino acid race-

mization, and radiometric techniques. Lucchitta (1979) drew upon this framework to link the integration of the lower Colorado River with the development of the Colorado River through Grand Canyon. The resulting paradigm, here called the Metzger model, represents a synthesis of many of the concepts introduced by previous investigators, invoking three major aggradation intervals and three major degradation intervals. The Metzger model (Fig. 2B), coupled with previously published geochronologic data and interpretations, can be summarized into the following stages: 1. Subsidence of a trough along the course of the lower Colorado River and incursion of an arm of the developing Gulf of California in the late Miocene to early Pliocene resulted in deposition of the Bouse Formation (Metzger, 1968). 2. Integration of the Colorado River sometime after 6.0 Ma (Metzger, 1968; Damon et al., 1978); regional uplift combined with deposition associated with the arrival of the river forced the arm of the sea back to the south (Lucchitta, 1979). 3. Canyon cutting followed by massive backfilling of sand and gravel sometime in the early Pliocene through early to middle Pleistocene time (between ca. 4.4 Ma and 700 ka) (Damon et al., 1978; Kukla, 1975); Unit B, first major aggradation. 4. Deep incision into the previous valley-filling deposits. 5. Backfilling with predominantly fine-grained sediments during the middle or late Pleistocene (the Chemehuevi Formation; ca. 700 ka to 35 ka) (Kukla, 1975; Bell et al., 1978; Blair, 1996; Lundstrom et al., 2000, 2004); Unit D, second major aggradation. 6. Progressive incision into the Chemehuevi fill interrupted by periods of stability and stream-terrace formation at discrete elevations above the river (one prominent level was dated as ca. 80 ka; Ku et al. [1979]). 7. Backfilling with channel and floodplain deposits to form the modern river environment (under way by ca. 8 ka); third major aggradation (Metzger et al., 1973). AN EVOLVING MODEL OF THE LATE CENOZOIC HISTORY OF THE LOWER COLORADO RIVER The existing data and concepts regarding the history of the lower Colorado River described above in the context of the Metzger model can be combined with recently identified geologic relations seen on this trip to formulate a more detailed working model of river inception and evolution over the past 5–6 m.y. (Figs. 2C and 2D). River Inception—By Lake or by Sea? The means by which the lower Colorado River developed its modern course remains the subject of debate. There are two contrasting models: (1) subsidence linked with early rifting in the Gulf of California extending up the lower Colorado River valley,

Birth of the lower Colorado River driving headward erosion that led to the progressive capture of upstream drainage systems (e.g., Lucchitta, 1972, 1979, 1998; Buising, 1990; Lucchitta et al., 2001); and (2) cascading lakespillover driven by upstream controls (e.g., Spencer and Patchett, 1997; Meek and Douglass, 2001). General versions of these two contrasting models were initially postulated by Blackwelder (1934) when he compared the dual possibilities that the river carved its valleys in response to regional uplift from near sea level or his preferred interpretation that the river formed via cascading lake-spillover. Geologic relations described in this guide are most consistent with the lacustrine spillover model. The crux of both models lies in their interpretation of the depositional environment of the Pliocene Bouse Formation (Metzger, 1968)—an enigmatic stratigraphic unit restricted largely to the river corridor and almost certainly linked with the river’s early inception (Lucchitta, 1972, 1979; Buising, 1990). The Bouse Formation typically has a thin basal deposit of marl or limestone overlain by varying amounts of mud, sand, and minor gravel, and by tufa lining the valley walls. It has been found in the string of basins along the river from Cottonwood Valley to downstream of Yuma (Metzger, 1968; Irelan et al., 1973). Its highest remnants are at ~330 m (~1082 ft) in the reach from Lake Havasu southward to the Chocolate Mountains of California and are well below sea level south of Yuma. In Mohave and Cottonwood Valleys, Bouse deposits have been documented as high as 550 m (1804 ft) in several locales. Bouse deposits are commonly thin where exposed along valley margins, but have been inferred to be several hundred meters thick in the subsurface in several basins along the lower Colorado River (Metzger et al., 1973; Metzger and Loeltz, 1973). At some locations Bouse deposits are interbedded with locally derived fanglomerates and early Colorado River deposits (Metzger, 1968; Dickey et al., 1980; Buising, 1990). The age of the Bouse Formation was initially estimated to be early to late Pliocene (Metzger, 1968; Carr, 1991). From results of our recent investigations, we can constrain its maximum age in Cottonwood and Mohave valleys to <5.5 Ma based on geochemical correlation of an underlying tephra to the 5.51 ± 0.13 Ma Connant Creek ash bed from the Heise Volcanic Field (Morgan and McIntosh, 2005).1 An exclusively marine-estuarine interpretation of the Bouse Formation, as suggested by paleontologic evidence (Metzger, 1968; Lucchitta et al., 2001), requires that a marine embayment extended from south of Yuma into Cottonwood Valley. Bouse outcrops as high as 550 m elevation near its northernmost known extent in Cottonwood Valley require at least that much regional uplift in the past 5.5 m.y. (Lucchitta, 1979). In this model, regional subsidence that permitted a marine intrusion up the future course of the lower Colorado River also drove headward erosion of a regional drainage system upstream, which resulted in eventual capture of a proto–Colorado River near the western edge of the 1

Previously, we have linked this bed to the related, but slightly older Wolverine Creek bed (e.g., various abstracts), but the geochemical distinction is difficult to make because the compositions are essentially identical.

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Colorado Plateau (Lucchitta, 1979; Lucchitta et al., 2001). The deep incision along the lower Colorado River after drainage integration is interpreted as a long-term response to regional uplift. Some of the difficulties posed to this model will be explored during this field trip. They include the following: 1. Headward erosion by a local or subregional drainage hypothetically proceeded rapidly from the upper end of the proposed marine embayment in Cottonwood Valley into the Colorado Plateau and breached several bedrock divides as it worked its way to the north and east, all without significantly impacting or exploiting tributary and closed basins immediately to the east and west of the lower Colorado River (e.g., Spencer and Pearthree, 2001). 2. The lowest portions of Mohave Valley must have subsided well below sea level by the end of the Miocene to allow the sea to transgress, only to rise in the past 5 m.y. to its current elevation, all in the apparent absence of concurrent major normal faulting. 3. The maximum elevations of the Bouse Formation do not gradually increase to the north, rather they increase abruptly north of Yuma and again at Topock Gorge (Spencer et al., 2005). 4. Strontium isotope ratios in Bouse deposits are more similar to modern Colorado River values than to marine values (Spencer and Patchett, 1997; Poulson and John, 2003). 5. New age controls indicate that the postulated marine incursion, river inception, river aggradation, uplift, and substantial valley carving would have to have transpired in a narrow window of time between 5.5 and ca. 3.3 Ma. 6. Major aggradation along the lower Colorado River postdates Bouse deposition and had to have occurred concurrently with the hypothesized regional uplift. Valleys were sites of major aggradation, not downcutting, immediately after the arrival of the river. The cascading lake-spillover model invokes headwater processes and requires no regional subsidence or uplift. In this model, some type of drainage rearrangement along the ancestral upper Colorado, possibly overflow from a terminal basin on the Colorado Plateau (e.g., Scarborough, 2001), ultimately resulted in divide-breaching in the eastern Grand Canyon area in the late Miocene. The Colorado River then developed a course through what is now the Grand Canyon and spilled into Grand Wash Trough and the Lake Mead area. Eventually, a divide in the Hoover Dam area was overtopped and the incipient Colorado River spilled through the series of basins along its modern course. The Bouse deposits found along the lower Colorado River thus may record a series of relatively deep, short-lived lakes fed by the earliest arrival of Colorado River water into the region. This model is bolstered by strontium isotope ratios in sediments and shells from the Bouse Formation that are consistent with lacustrine basins fed by the Colorado River and are inconsistent with marine influence (Spencer and Patchett, 1997; Gross et al., 2001). The basic premise of the lake-spillover model has been criticized by proponents of the marine incursion hypothesis

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largely on the basis of paleontologic evidence from the Parker area southward (e.g., Lucchitta et al., 2001), and the possibility of postdepositional alteration of carbonate Sr isotope ratios (Lucchitta, 1998). Until recently, however, no unequivocal stratigraphic evidence had been found along the lower Colorado River to strongly support either mode of river inception. This field trip guide describes sites on both sides of the paleodivide between Cottonwood and Mohave valleys that exhibit stratigraphic evidence consistent with the lake-spillover model (Figs. 2C and 2D; Faulds et al., 2004; House et al., 2004b). In the northern Mohave Valley, a coarse fluvial conglomerate derived from a nearby, upstream source (the gravel of Pyramid hills) unconformably overlies a thick-sequence local Miocene fanglomerate and axial channel gravels and, in turn, is conformably overlain by the Bouse Formation (Fig. 2C). The foregoing sequence is, in turn, unconformably overlain and deeply incised by an intricately bedded fluvial deposit of the early Colorado River. Its base (the gravel of Panda gulch), is less than 20 m above modern river level and forms the lowest part of a thick deposit of Colorado River alluvium (alluvium of Bullhead City) that is interbedded with tributary fan gravel ~230 m higher above the river. The composite section reflects a sequence of a local basin conveying a large flood, followed by deep inundation, followed by development of a major through-going drainage. The stratigraphic sequence in southern Cottonwood Valley closely parallels and complements the Mohave Valley record (Fig. 2D). There, the late Miocene fanglomerate sequence grades upward into a conspicuous sequence of flat-bedded mudstone and sandstone (the Lost Cabin beds) with characteristics of lacustrine and low-energy fluvial deposition. A minor unconformity at the top of the fine-grained Lost Cabin beds is overlain by the Bouse Formation, which is, in turn, overlain along a major unconformity by the deposits of the early Colorado River (alluvium of Bullhead City). This composite record reflects a sequence of two phases of inundation of Cottonwood Valley followed by the development of a major, through-going drainage. We suspect that the first phase of inundation terminated with the flood associated with the gravel of Pyramid hills (lake-spillover) and that the second phase of inundation involved the deposition of the Bouse Formation in both valleys. Our recent mapping efforts have also contributed to the clarification of the chronology of this sequence of events. Tephra deposits found in fanglomerate in Mohave Valley and in the Lost Cabin beds in Cottonwood Valley constrain the age of Bouse deposition to <5.5 Ma. We discovered a 4.2–3.6 Ma tephra bed in the upper 30 m of the alluvium of Bullhead City on the piedmont of the Black Mountains (Faulds et al., 2002). This tephra indicates a maximum age for the culmination of the primary river aggradation phase. We also found a 3.3 Ma tephra bed in younger alluvial deposits that truncate the upper part of the thick river fill which indicates that the river had begun to incise by this time. These temporal constraints on river inception and early evolution are similar to those developed for the Colorado River at the mouth of Grand Canyon described previously, and the overlap

strongly suggests a linkage between canyon incision upstream and thick aggradation in valleys downstream. Pleistocene History of the Lower Colorado River in the Field Trip Area The bed of the lower Colorado River through Cottonwood and Mohave valleys was roughly 250 m higher than the modern river by the middle Pliocene. The river subsequently downcut several hundred meters to near the pre-integration valley elevation. A tremendous amount of sediment must have been removed during this period, but the duration of the episode is not well constrained. Suites of progressively lower terrace remnants and coeval alluvial fans that exhibit strong carbonate soils (Stages V+ and VI) suggest that much of deep incision may have transpired in the late Pliocene into the early Pleistocene. Our recent mapping in the Laughlin area has revealed evidence for at least 3 Pleistocene aggradation events since incision to the pre-integration level. The first deposits laid down by the river following incision into the Bullhead sediments include the Riverside beds and the conglomerate of Laughlin. The Riverside beds comprise an interbedded sequence of mud, sand, and gravel. They are separated from a subsequent series of similar river deposits by the conglomerate of Laughlin, a coarse conglomerate from a postincision flood that further gouged out part of the Pyramid hills. It is composed of a mixture of far-traveled river gravel and large gravel clasts of locally derived granite. A paleochannel formed by the flood has subsequently been backfilled by a thick deposit of Pleistocene river alluvium. This subsequent package (described below) also overlies erosional topography and a moderately developed paleosol (Stage III carbonate) in the underlying deposits. The Chemehuevi Beds A series of conspicuous alluvial fills exist along the river between the mouth of the Grand Canyon and Yuma. They are typically characterized by a relatively thick basal unit of flat-bedded mud and fine sand overlain by a comparably thick unit of loose, gravelly sand, often forming conspicuous piles of sediment strewn over a rugged bedrock substrate. Newberry (in Ives, 1861) and Lee (1908) concluded that this package was the result of river aggradation. Longwell (1963) concluded that it was a deltaic sequence deposited in a lake(s) associated with valley blockage at unknown downstream locations. Metzger et al. (1973) concluded that the deposits were associated with fluvial aggradation linked to climate variability. The name Chemehuevi Formation has been in common use and now has a general connotation for all fill sequences along the lower Colorado River from Lake Mead to Yuma characterized by flat-bedded mud and fine sand overlain by a looser sequence of gravelly sand beds or of fluvial gravel. Our studies suggest that this type of sequence is not unique in the history of the river and that the Chemehuevi Formation as envisaged by Longwell (1936) may contain a series of similar, disconformable sequences. Also, the upper sandy package in the archetypical sequence (as envisaged by Longwell, for example) typically overlies the fine-grained package along an erosional

Birth of the lower Colorado River unconformity. This is one of the reasons that Metzger et al. (1973) opted to drop the Chemehuevi Formation name altogether in their studies. Instead, they envisioned a lower unit of mud and fine sand, which they called unit D, and an overlying unit of gravelly sand that they called unit E. In this report, we informally refer to the sequence as the Chemehuevi beds for reasons outlined below. All investigators have agreed that at least one major aggradation event is recorded by the Chemehuevi deposits. Longwell (1936) reported no unconformities in the sequence and concluded that it was the result of a single, thick, and largely uninterrupted aggradation event. In this model, all subsequent fluvial landforms recorded only intermittent hiatuses during net incision into the single fill and the only notable subsequent aggradation is associated with the Holocene fill below the modern floodplain (Metzger et al., 1973). Two recent studies focused on an inferred anomalous sedimentology of the Chemehuevi beds within the context of an uninterrupted episode of aggradation. Blair (1996) interpreted the Chemehuevi Formation as evidence for an abrupt change in the volume and rate of floodplain sedimentation, likely owing to a climatic perturbation. Lundstrom et al. (2000) invoked a singleaggradation model and speculated that the entire fine-grained sequence may have been deposited by a single flood. More recently, Lundstrom et al. (2004) suggested several plausible mechanisms. In a series of geologic maps, Faulds (1996a, 1996b) broadly classified Chemehuevi deposits within an undivided unit of river alluvium spanning the Pliocene and Pleistocene. We have recently mapped a similarly distinctive deposit of predominantly flat-bedded, fine-grained fluvial mud and sand overlain unconformably by a sequence of loose medium sand to gravelly sand in the Laughlin area (Faulds et al., 2004). We have also identified a prominent unconformity within the lower, fine-grained part of this package that juxtaposes two very similar deposits of fluvial mud and fine sand. The overlying gravelly sand caps fluvial terraces that are roughly 100 m above the modern channel of the lower Colorado River. This entire package (upper sand and two lower mud deposits) overlies a paleosol and erosional topography in a similar set of fine-grained Colorado River deposits (the Riverside beds) and a conspicuous, coarse fluvial conglomerate (conglomerate of Laughlin). Thus, the Chemehuevi Formation of Longwell (1936, 1963) is locally, at least, an assemblage of a series of disconformable packages of river laid alluvium that have a generally homogeneous appearance. For this reason, we use the broader designation of the Chemehuevi beds to distinguish the composite package as a single map unit. Division of the Chemehuevi beds into a series of upper and lower components as well as a fluvial gravel component provides for a more detailed characterization of the deposit where warranted. The Mohave Sediments The Mohave sediments comprise a package of fluvial sand, mud, and gravel that underlies the terrace below the Mohave Generating Station in Laughlin (the Mohave terrace). The terrace surface is ~60 m above the modern lower Colorado River channel. Deposits below the terrace surface share compositional similarities

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with the Chemehuevi beds exposed elsewhere, but their stratigraphic context suggests that they may be a younger sequence. Compared to the exposures of Chemehuevi beds a mile to the north, this package contains a more variable assemblage of fluvial mud and sand overlying a conspicuous coarse-grained channel gravel and interbedded with tributary fan gravels. The core of the fluvial package interfingers toward the west with alluvial fan deposits derived from the Newberry Mountains and to the east with alluvial fan deposits from the Black Mountains; thus, tributary deposits constrain a paleochannel position beneath the Mohave Generating Station (the Mohave paleochannel). The bounding fan deposits grade upward into a widespread piedmont map unit—Qai2. That relation suggests this sequence is younger than deposits exposed in the Davis Dam bluffs, where the upper Chemehuevi beds are coeval with an older piedmont alluvial unit (Qai1). We have designated the Mohave paleochannel deposits the Mohave sediments and consider them the youngest deposits in the Chemehuevi beds. Without the bounding tributary deposits it is difficult to confidently identify Mohave sediments as a discrete package. Age of the Chemehuevi Beds and Similar Packages The ages of the Quaternary alluvium have been harder to pin down than the ages of the Miocene to Pliocene deposits. To date, the existing data are contradictory and compromised by incomplete or ambiguous exposures of stratigraphy. However, there are several interesting lines of evidence. Fossils. Newberry (in Ives, 1861) reported the first Pleistocene fossil discovery along the river in upper Cottonwood Valley when he extracted a mammoth tooth from lower Chemehuevi muds. The species and age of the specimen are uncertain. Metzger et al. (1973) discovered a Pleistocene mammoth tusk in the Chemehuevi beds. Bell et al. (1978) performed U-series analysis on the specimen and reported an age of ca. 102 ka. In contrast, an amino acid racemization analysis on a mammoth skull from lower Chemehuevi-like beds near Parker, Arizona, yielded an age of 900 ka, and nine small vertebrate fossils from near Blythe, California, yielded amino acid ages of ca. 100 ka (Lee and Bell, 1975). According to L.D. Agenbroad (1995, personal commun.), there is evidence for 3 Mammuthus meridionalis fossils along the river between Laughlin to just south of Parker, and their most likely age range is 1.5–1.7 Ma. Isotopic and Luminescence ages. Samples of intact and unaltered wood are rare in the Chemehuevi beds. However, in one instance, Blair (1996) collected a wood specimen from mud in the lower Chemehuevi beds near the north end of Cottonwood Valley, which yielded a radiocarbon age of 35 ka. Lundstrom et al. (2004) reported that a series of luminescence ages from lower Chemehuevi deposits from various sites along the river range from ca. 40 ka to 70 ka. They also obtained U-series ages on carbonate-coated gravels from inset younger terrace surfaces that ranged in age from 32 to 60 ka. Paleomagnetic data. Paleomagnetic analyses of Chemehuevi deposits from near Parker, Arizona, showed that they are normally magnetized, and thus less than 780 ka (Bell et al., 1978).

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Lundstrom et al. (2000, 2004) also report normal magnetization in the Chemehuevi beds from various sites. Tephra beds. Longwell (1936) reported on and sampled a 3 m thick tephra bed at the base of his Chemehuevis Formation near the town of Callville, Nevada (now submerged by Lake Mead). Neither the source nor the age of this ash is known. The Lava Creek B ash (0.62 Ma) and either the Bishop (0.74 Ma) or the older Glass Mountain ashes are Pleistocene ashes previously reported from the general area of the lower Colorado River (Whitney, 1996; Merriam and Bischoff, 1975); A. Sarna-Wojcicki, 1998, personal commun.), but have not been collected from Colorado River deposits. Recently we (Brock, House, and Pearthree) discovered and sampled a tephra bed from the lower Chemehuevi–type muds north of Katherine’s landing. The stratigraphic context of the bed indicates to us that it is of Pleistocene age. Analysis and identification of this tephra by the U.S. Geological Survey is pending. The Pleistocene stratigraphy in the Laughlin area indicates a complex history of processes along the river during that time. There is evidence for at least three likely Pleistocene-age sequences of predominantly fine-grained Colorado River alluvium separated by unconformities. This record is difficult to resolve from the basis of a series of similar but scattered outcrops of alluvium, and is inferred from detailed study in a relatively small area. We believe that the Chemehuevi beds may represent a series of responses of the fully integrated Colorado River to a cyclic, external control—presumably, but not certainly, climate.

5. Incision through the Bouse Formation and underlying deposits due to lake spillover and drainage through Topock Gorge; arrival of bedload associated with the Colorado River (gravel of Panda Gulch). 6. At least 250 m of river aggradation culminates between 3.6 and 4 Ma (alluvium of Bullhead City). 7. Deep incision begins by 3.3 Ma and eventually reaches near modern river level (series of fans and river terraces with strong carbonate soils). 8. Minor aggradation ca. 1.5–1.7 Ma. (Riverside beds). 9. Locally catastrophic flood (conglomerate of Laughlin). 10. Multiple aggradation/degradation cycles up to 35 ka (lower Chemehuevi beds, upper Chemehuevi beds, Mohave sediments). 11. Progressive incision, strath formation, minor aggradation (alluvium of Emerald River, alluvium of Big Bend, various strath veneers). 12. Incision followed by Holocene aggradation to form the modern floodplain (Metzger et al., 1973; alluvium of Riviera). FIELD TRIP Day 1: Las Vegas to Laughlin: The Debut of the Colorado River Total driving distance: ~100–120 mi (Figs. 1 and 4).

FINALLY—OUR WORKING MODEL The working model that we have developed for the late Tertiary and Quaternary record of the lower Colorado River based on relations in the Laughlin area is outlined below and shown schematically in Figure 3. Age estimates for events in the Quaternary are not closely constrained at this time. Parenthetical terms (below) are informal names that we have developed for representative stratigraphic units in the Laughlin area that record the sequential events, including a variety of late Pleistocene terraces and strath veneers that are not described in the previous sections. The chronological assignments are based on new data and our interpretation of the likely context of previously reported data and are, in some cases, conjectural placeholders. 1. Local basins filling with extensive alluvial fans (Miocene fanglomerate). 2. Transition from local fan and axial channel deposition to intermittent lacustrine conditions in Cottonwood Valley prior to 5.5 Ma (Lost Cabin beds); roughly concomitant development of axial drainages in northern Mohave Valley (axial gravel complex). 3. Spillover/divide-failure through the Pyramid hills and major flooding from Cottonwood Valley into Mohave Valley sometime after 5.5 Ma (gravel of Pyramid hills). 4. Deep lake formation in Mohave Valley and Cottonwood valleys soon after 5.5 Ma owing to damming at Topock divide (Bouse Formation).

Directions and Highlights en Route to Stop 1.1 The field trip begins at the intersection of Hwy 93 and Hwy 95 between Henderson and Boulder City, Nevada (0.0 mi) (Fig. 1). From here, drive south on Hwy 95 toward Laughlin, Nevada. At about the 10 mi point, the route passes the turnoff to Nelson’s Landing, Nevada, the site of a catastrophic and fatal flash flood down Eldorado Canyon on 14 September 1974 (Glancy and Harmsen, 1975). Enter Searchlight, Nevada, at 35.7 mi atop the 3500-ft divide between Eldorado Valley (closed basin) and Piute Valley (lower Colorado River tributary). About 20 mi south of Searchlight, exit left (east) on Hwy 163 to Laughlin, Nevada (reset odometer). This route traverses a pediment on the west flank and in the interior of the rugged Newberry Mountains, underlain mainly by Miocene granite. Note the changing regional topography as you continue toward Laughlin. The floor of Piute Valley is at ~2516 ft (767 m), and the highway quickly climbs to 2960 ft (901 m) before it begins a steep descent to the Colorado River at 500 ft (152 m). The highway eventually debouches onto the steeply sloping head of the Dripping Springs Wash alluvial fan at the foot of the Newberry Mountains. At ~2.5 mi below the mountain front, the highway flattens slightly as it skirts the southwestern edge of the Davis Dam terrace, the higher of two major Pleistocene river terraces in the Laughlin area. To the southeast, the Mohave Generating Station sits atop the other, lower terrace which we call the Mohave terrace. Exit right on Civic Drive near the base

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Extremely strong soil (Carbonate Stage V+ to VI) 1320 v-v-v-v-v-v-v-v-v-v-v-v

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Colorado River 500' (152 m)

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Figure 3. Schematic representation of late Cenozoic stratigraphy of the Colorado River and its piedmont tributaries in northern Mohave Valley. Nomenclature described in text.

of the piedmont slope and then turn right (south) onto Edison Drive (Fig. 5). Large boulders exposed here are part of the Laughlin conglomerate. The coarse conglomerate is unconformably overlain by the Chemehuevi beds that are visible along the north side of Edison Blvd. Follow Edison Blvd. to the top of the Mohave terrace across an array of late Cenozoic river and alluvial fan deposits. Crest the terrace and take a right turn at the traffic signal (1.95 mi) onto Casino Drive (southbound). Pass the Harrah’s Casino and continue until you reach a deeply incised wash visible along the east side of the road (~0.5 mi from traffic light). Park on the east side of the road just above the wash (see Fig. 5 for specific locations). Stop 1.1: The Laughlin Bluffs A series of steep gulches carved in the east face of the Laughlin bluffs expose the key stratigraphic evidence that documents a

change from local, possibly enclosed drainage, to deep inundation with standing water, and then finally to a major throughgoing river. We refer to the deposits recording these events as the transitional sequence. The sequence here is linked to a related sequence of concurrent and complementary changes in depositional conditions in Cottonwood Valley (see Day 3). The deposits of the transitional sequence form a spine of relatively indurated late Miocene to early Pliocene sediments surrounded and overlain by various Quaternary alluvial fan and Colorado River deposits resulting in a complex assemblage of deposits (Fig. 6). Stop 1.1a: Panda gulch. The first stop will focus on exposures of the transitional sequence in Panda gulch. Each member of the sequence is clearly visible on the northwest facing slope above the gulch (Fig. 7A). Late Miocene, untilted fanglomerates from the Newberry Mountains form the base of the section and are overlain by a cross-stratified channel gravel deposit contain-

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Figure 4. Map showing physiographic setting and location of field trip sites in Southern Cottonwood Valley and Mohave Valley.

ing clasts of local rock types—the “axial gravel” (Tag; Fig. 7B). We interpret these gravels as marking the position of an axial channel near the head of Mohave Valley at a time when it was not directly connected with drainage through Cottonwood Valley. The axial gravel and underlying fanglomerate are unconformably overlain by the gravel of Pyramid hills (Tpg; informally named Pyramid gravel; Fig. 7C), a conspicuous, immature boulder conglomerate. Deposition of the Pyramid gravel was accompanied by erosive enlargement of the axial channels and locally deep scour into the underlying fanglomerate. The

Pyramid gravel is predominantly a cobble-boulder conglomerate with a maximum thickness of ~20 m. Many clasts are subrounded to rounded. Some exposures contain sparse, locally derived sediments (reworked fanglomerate), and others are dominated by thick, stratified sequences of monomictic grussy sand and subrounded pebble-gravel possibly derived from regolith stripped from slopes of the Pyramid hills. Cross-stratification is evident in some intervals of the Tpg (e.g., trough crossbedding; Fig. 7C). Many exposures show a clast-supported structure, whereas some are matrix-supported and slurry-like. Most outcrops of the Pyramid gravel are composed almost entirely of cobbles and boulders of megacrystic granite from the Pyramid hills and reworked local fanglomerate from the Newberry Mountains. We interpret the Pyramid unit as a catastrophic flood deposit from a clear-water breach through a paleodivide in the Pyramid hills. Deposits in Cottonwood Valley (Stop 3.1) indicate a roughly concurrent lacustrine environment some 12 mi upstream of the Pyramid Hills. The Pyramid gravel is conformably overlain by the Bouse Formation, indicating inundation of this area by standing water following the deposition of the gravel (Fig. 7A and 7D). Outcrops

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Figure 6. Detailed geologic map of the Laughlin bluffs area (modified from Faulds et al., 2004). See text for discussion of various units. Tfn—Miocene fanglomerate from Newberry Mountains (Tfn1—preBouse; Tfn2—syn/post-Bouse); Tag—axial gravel; Tpg—Pyramid gravel; Tb—Bouse Formation; Tbhl—lower Bullhead alluvium (Panda gravel); Qcm—Mohave sediments (Qcmf—fine facies; Qcmg—gravelly facies); Qai—intermediate age alluvial fans (late Pleistocene?); Qer—alluvium of Emerald River; Qay—young alluvial fans (Holocene); Qcr—alluvium of Riviera; Qx—extensively disturbed areas.

of Bouse in the Laughlin bluffs range from 0.5 to ~3 m thick and include beds of marl, mud, and minor sand. The outcrops in the Laughlin bluffs occur only 140 ft above the surface of the modern Colorado River. The sequence of Bouse over Pyramid gravel suggests that a catastrophic flood from an upstream source, possibly a lake, was immediately followed by quiescent deposition in a large body of standing water.

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The preceding sequence is unconformably overlain and deeply incised by gravels and sands of the early Colorado River—the alluvium of Bullhead City (Tbh; informally called the Bullhead alluvium). The Bullhead alluvium is generally equivalent to units A and B of Metzger et al. (1973). The gravel at the base of the Bullhead unit is largely comprised of locally derived sand and gravel reworked from Newberry Mountain–sourced fanglomerates, but is also peppered with well-rounded pebbles of chert and well-rounded cobbles of diverse lithologies from far upstream. The mix of light colored local sand and gravel with the exotic, mostly dark-colored pebble component imparts a distinctive black and white speckled appearance to the unit, hence our informal name Panda gravel (Tbhl; Fig. 7E). The Panda gravel has a cobble-rich basal conglomerate containing mainly clasts of reworked fanglomerate and Pyramid gravel. Look on the slopes for examples of fluted (stream-worn) boulders. The base of the Panda unit dives down-section into the late Miocene fanglomerate down the wash (Fig. 8) and then climbs up toward the east and the modern course of the Colorado River. The base of the Panda gravel here thus defines a large paleochannel. Evidently, the arrival of the river in northern Mohave Valley immediately preceded or was accompanied by an interval of erosion following the recession of the Bouse water body. The step-like geometry in the Panda gravel channel here (Fig. 8), the occurrence of stacks of laterally continuous layers of coarse gravel, and the abundance of locally derived clasts in the body of the deposit suggests lateral erosion and reworking of the Miocene substrate by a rapidly aggrading, greatly over-fit river. The Bullhead unit is discontinuously exposed from this point (~20 m above the river) up to gravel lags and river sands interfingered with Black Mountain alluvial fan gravels at levels as high as 250 m (~820 ft) above the river. Stop 1.1b: Lavender gulch. A thicker section of the Pyramid gravel (similar to that shown in Fig. 7C) can be evaluated in Lavender gulch ~0.5 mi to the south of Panda gulch. In general, the deposits continue to thicken toward the south from here until the Big Bend of the Colorado, beyond which the transitional section is not preserved. To reach this site, drive south on Casino Drive from the Panda gulch pullout for ~0.45 mi. Once you reach the surface of the Mohave terrace, there is an expansive flat area along the east side of the road. Park there. Look for a narrow terrace ridge with a clear vehicle track that extends quite far toward the east. Head down the steep north side of the ridge. About halfway down this traverse, a 3–5 m thick interval of the Bouse Formation is exposed. The gulch terminates in a precipitous pour-over at the contact between the Pyramid gravel and the axial gravel. An excellent perspective on the obtrusive nature of the Pyramid gravel, its channel forms, and its erosive contact is also available from Arizona Hwy 95 in Bullhead City. The Pyramid unit forms a prominent dark swath in the middle of the Laughlin bluffs and many paleochannel forms are easily discerned. Directions and Highlights en Route to Stop 1.2 Return to Casino Drive and continue south. The road turns westward along the Big Bend of the Colorado River (Fig. 5). Deep

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A

Figure 7. Annotated photographs of the late Miocene–early Pliocene transitional sequence and bracketing deposits exposed in upper Panda Gulch and key locations in the Laughlin Bluffs. (A) Complete, compressed section. Tfn—Newberry Mountains–derived fanglomerate; Tag—axial gravel; Tpg—gravel of Pyramid hills; Tb—Bouse Formation; Tbhl— lower alluvium of Bullhead City (gravel of Panda gulch); Qcm—Quaternary Colorado River deposits (Mohave sediments). (B) Paleochannel in axial gravel unit (Tag). (C) Gravel of Pyramid hills in channel carved in fanglomerate. Tpg is ~20 m thick here. (D) Bouse Formation limestone overlain by the gravel of Panda Gulch. (E) Base of paleochannel incised in fanglomerate and filled with the gravel of Panda Gulch.

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washes incised in the Mohave terrace to the north expose some of the transitional sequence and parts of the Mohave paleochannel, a much younger package of Colorado River sediments, and interbedded alluvial fan gravels are exposed in Arizona gulch, the largest drainage on the Mohave terrace (Fig. 6). Fan gravels exposed in roadcuts between Panda gulch and the major bend in the highway at ~0.5 mi are derived from the Black Mountains and bound the eastern margin of the paleochannel. Heading due west, outcrops to the north show late Pleistocene alluvial fans derived from the Newberry Mountains and interbedded Colorado River sediments that define the western margin of the paleochannel. The road then drops onto the toe of the Hiko Springs alluvial fan, the terminus of one of the largest drainages on the Newberry Piedmont and site of the tallest flood control structure in Clark County (2.5 mi up the wash). At the stoplight, turn south (left) on the Needles Hwy (4.8 mi from the Edison Drive–Casino Drive intersection; Fig. 9). The road crosses an unnamed wash and drops onto the historical (predam) floodplain of the Colorado River. On the west side of the road is an exposure of the Newberry detachment fault. The reddish rock clinging to the base of the Newberry Mountains is an upperplate remnant of Proterozoic granite. Beyond the fault exposure, the road continues to follow the historical floodplain and then climbs through a sequence of late Miocene, pre-Bouse fanglomerate. At just ~2 mi from the last stoplight, look for an unmarked exit on your right to reach a frontage road paralleling the west side of the highway. Continue south on the frontage road. After 2.3 mi, turn right (W) on the pipeline road and proceed up the piedmont between the Newberry and Dead mountains. Much of the piedmont is a pediment formed on late Miocene fanglomerate and granite. The fanglomerates extend to near modern river level and are likely coeval with fanglomerates in the Laughlin bluffs. Outcrops of the Bouse Formation overlying the fanglomerate are exposed 0.5 mi due west from this intersection. Follow the pipeline road for 2.3 mi. Turn left (S) at the intersection with a powerline road. Follow this road for ~1 mi and turn left (E) down a narrower track. Follow this track for 0.4 mi to the bedrock outcrops. Stop 1.2: Manchester Beach This stop highlights a variety of deposits of the Bouse Formation that illustrate the extent of the body of standing water

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Tbhl

Tag Tfn

Tfn

Tbhl

Figure 8. Annotated photo mosaic of exposure in Panda gulch showing the disconformable basal contact of the Panda gravel (Tbhl) over axial gravels (Tag) and fanglomerate (Tfn).

that formed in Mohave Valley following the large flood through the Pyramid hills divide. An extensively pedimented bedrock outcrop on the northeast edge of the main block of the Dead Mountains is underlain by Miocene granite and quartz diorite (House et al., 2004a). Along the northern margin of the pediment at this stop lie a series of enigmatic deposits of locally derived, rounded gravels. We interpret these as beach gravels deposited along the margin of the Bouse basin. The gravels are locally interbedded with sandstone and calcareous mud that onlap the bedrock. They occur as relatively flat benches hemmed in by bedrock protrusions along the northern margin of the pediment. They have a distinctive yellowish-orange oxidized patina. The lag

of rounded, local gravels can be traced north across the piedmont to a sparse lag of rounded gravels on deeply weathered, craggy fanglomerate remnants and to stratigraphic exposures at ~548–560 m (1800–1840 ft) in washes draining the Newberry Mountains. Some of the gravel beds have unidirectional crossstratification and are locally interfingered with thin beds of flatlying sandstone. The distribution of various deposits of the Bouse Formation in this part of Mohave Valley are suggestive of a “mega-drape” of sediment that covered a preexisting, deep valley similar to the present one (Fig. 10). Bouse sediments on the north flank of the Dead Mountains pediment occur from 1320 to 1740 ft elevation. A marl outcrop along the powerline road overlies fanglomerate

ins way Bouse outcrops

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Figure 9. Air-photo map of route to and location of Stop 1.2.

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along a span from 760 ft to 1000 ft elevation. Five mi (8 km) south of the Manchester beach site, Bouse roundstone gravels are found in association with tufa encrustations on the bedrock face of the Dead Mountains and tufa-cemented colluvium between 350–440 m (1160 and 1440 ft) above sea level (asl). We have also found large tufa clasts on late Miocene fanglomerate ~1.5 mi (~2.4 km) southeast of here at 365 m (~1200 ft). The tufa rinds and beach gravels at successively lower elevations probably indicate periods of stasis in a fluctuating (possibly slowly draining) deep body of water. Additional Bouse outcrops along the Dead Mountain piedmont include beds of marl that steeply onlap bedrock at ~305 m (~1000 ft), and project below thick deposits of mud and sand that are exposed between 170 and 305 m (560 and 1000 ft). In summary, the distribution of Bouse exposures in this part of Mohave Valley ranges from 170 m (560 ft) to at least 548 m (1840 ft). In most cases, Bouse sediments overlie late Miocene fanglomerate or sit directly on bedrock. The highest clastic deposits that we have identified in this area may approximate the peak elevation of the Bouse water body in Mohave and Cottonwood valleys. They occur at the same elevation as the highest previously documented Mohave Valley Bouse outcrops in Silver Creek Canyon (Stop 1.3) (Metzger et al., 1973; Spencer and Patchett, 1997), and they also occur at the same general elevation of the highest outcrops of likely Bouse basin margin deposits that we have found in Cottonwood Valley (Stop 3.1), 18 mi (30 km) to the north. The Black Mountains piedmont. The next several field trip stops are on the Black Mountains piedmont (Figs. 4 and 11) and focus on evidence for the timing and nature of maximum valley inundation during the Bouse interval and the subsequent major aggradation and incision of the early Colorado River. The character and distribution of deposits preserved on the Black Mountains piedmont provide insights into the timing, nature, and duration of the Bouse interval, excellent evidence for the maximum aggradation of the Colorado River to 250 m (820 ft) above the modern river, and evidence for the timing of initial incision of the river after maximum aggradation. Tephra beds discovered on this piedmont provide timing constraints for each of these important intervals, restricting all of this activity to the period between the Miocene–Pliocene boundary and the middle Pliocene (between ca. 5.5 and 3.3 Ma; Fig. 12). The southern Black Mountains consist of early to middle Miocene volcanic rocks and are lithologically distinct from the predominantly granitic rocks exposed on the west side of Mohave Valley. These rocks are much less tilted and extended than coeval rocks across the river. Reconstructing eastward slip on the east-dipping Newberry detachment fault across the river to the west (Spencer, 1985), the extrusive rocks in the Black Mountains would rest approximately above middle Tertiary granitic rocks in the Newberry and Dead Mountains 20–30 km to the west. The Black Mountains are capped by mesa-forming olivine basalt dated at 15.8 Ma, which unconformably overlies the early Miocene sequence (Gray et al., 1990) and is a major source of fan gravels on the piedmont.

Directions and Highlights en Route to Stop 1.3 Retrace your route back through Laughlin to Hwy. 163. Turn right (E) and cross the Colorado River on the Laughlin bridge (reset odometer). Continue straight through the stoplight immediately east of the river onto Bullhead Parkway. The road ascends the lower piedmont of the Black Mountains and bends to the south, passing Bullhead City Airport. At ~0.9 mi, there is a small exposure of the basal limestone of the Bouse Formation in a low roadcut on the east side of the road. Just farther to the east and below the Bouse outcrops there are several small outcrops of the Pyramid gravel. At ~2.2 mi, the route approaches higher ridges that extend down the deeply dissected piedmont. Many rounded ridges are 10–40 m higher than adjacent valley bottoms. The ridges are almost uniformly capped by several meters of tributary gravel with moderately to strongly developed petrocalcic soil horizons. The oldest remnants may date to the middle Pliocene (site 1.5). Many of the ridges beneath the capping gravel are composed of Colorado River deposits of the Bullhead alluvium (Fig. 13). Continuing to the south, there are several excellent exposures of weakly to moderately indurated, cross-bedded sand and rounded river gravel of the Bullhead unit. The elevation here is ~280 m (920 ft) asl, roughly in the middle of the major Pliocene river aggradation sequence. At 3.3 mi, turn left (E) across the northbound lane onto a dirt track in the valley of Secret Pass Wash, a large drainage that heads in the Black Mountains. Along the route in the active channel you will see as many as three prominent Pleistocene terraces in the valley. A few good exposures along the valley sides reveal predominantly river deposits with some tributary fan gravel. Between 1.2 and 1.4 mi up the wash there are better exposures of the Bullhead unit interfingered with tributary fan deposits on the north side of the valley, up to ~360 m (1180 ft) asl. This is near the upper limit of voluminous Colorado River deposits, although limited river sand and gravel deposits and lags can be found at numerous localities up to 400 m (1320 ft) asl in this area. The road forks at ~2.2 mi; stay in the valley bottom by bearing left. As the valley narrows and you will see several rotated blocks and pillars of indurated fanglomerate on the north side of the valley. These blocks have slid down over fine-grained deposits of the Bouse Formation, which is exposed at ~2.5 mi. Proceed to 3 mi and stop along the road. Stop 1.3: Secrets of Secret Pass At this stop the character of the Bouse Formation can be examined, including new maximum age constraints on its deposition. Here is an ~20 m thick exposure of the Bouse that has most of the components described by Metzger (1968) and Metzger and Loeltz (1973). The basal unit is a thin, poorly exposed limestone overlying fanglomerate at the north edge of the wash. This relationship is better exposed and higher above the wash upstream, and not exposed downstream, implying that the fan surface on which the Bouse Formation was deposited sloped more steeply to the west than the modern channel (~2° versus 1.5°). The basal limestone is overlain by pale green mudstone

Birth of the lower Colorado River

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approximate maximum Bouse water level Bouse gravel, sand, limestone, and tufa Bouse limestone, mud, and sand

12000

fan

glo

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ra

Pyramid gravel

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12000

5.5 Ma tephra

Newberry Mountains 800

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371 20000

Bouse marl, sand, and tufa

400

10,000

30,000

20,000

Horizontal Distance, West to East (ft)

Figure 10. Schematic cross-profile of Mohave Valley showing distribution of key late Neogene stratigraphic units. Positions of mapped outcrops of the Bouse Formation are shown in black. Inferred extent of Bouse shown in dark gray. Inferred sequence of axial channels also shown. Depiction of uppermost axial channel overlain by the gravel of Pyramid hills represents relations in Laughlin bluffs.

Tbh

Lower Nomlaki tephra Qa Tbh

Tbh

Qa

h

Secret PasTbhs Was Bullhead Parkway

Tbh

Tbh

Qa

Tbh

QTa

1320 ft asl Qls

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QTa

Tfb

QTa

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Tbo

Ttc

Stop 1.3

Tfb

Qa

QTa

1440 ft asl

Tfb

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Connant Creek tephra

Bouse outcrop

Tfb

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R

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1300 ft asl

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Nomlaki tephra 1300 ft

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asl

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Map Units

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QTa

Tbh Tbh

Qa

Qa - Quaternary alluvium QTa - middle Pliocene to Tbh earlyQTa Quaternary alluvium Tbh - early Pliocene Colorado Tfb QTa Tbh RiverQadeposits Tbf - latest Miocene to early Pliocene Black Mountain fanQa deposits

Stop 1.4

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QTa

ve

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QTa

Tfb

rC

ree

k

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Q

Black Mountains

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Bouse outcrop

1700 ft asl R

Tfb

R Q

1800 ft QTa asl

Stop 1.5

Figure 11. Surficial geologic map of part of the Black Mountains piedmont. The field trip route is shown by a heavy gray line and field trip stops are identified.

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maximum Bouse water level

1800

Tbo

Mio-Pliocene < 5.5 Ma

Altitude (ft)

1600

n fa e 1 en Tfb c io s M sit te o la dep

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Connant Cr tephra 5.5 Ma

800 1800

1800

Tbo

Middle Pliocene <4.2 Ma, >3.3 Ma 1600

Altitude (ft)

lower Nomlaki tephra 3.6-4.2 Ma

1400 max Colorado River aggradation

1600

fan ene lioc Tfb2 P ly ts ear eposi d

1400

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1000

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Altitude (ft)

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1400 Nomlaki tephra 3.3 Ma

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728000

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UTM Easting (m) Figure 12. Schematic cross sections of upper Black Mountains piedmont showing relationships at the time of the Bouse lake filling, maximum Colorado River aggradation, and the beginning of Colorado River incision after maximum aggradation.

Birth of the lower Colorado River that grades up into brown mudstone on the lowermost slopes north of the wash. The mudstone is overlain by more indurated tan sandstone and siltstone beds; many of the beds are massive, but locally the sandstones have cross-bedding. The uppermost sandstone and siltstone beds are interfingered with beds of fine, locally derived gravel. These Bouse sediments likely were deposited in a near-shore environment with abundant clastic sediment input of material from tributary drainages off of the Black Mountains. This site is more than 100 m below the maximum level of inundation, but it is not clear whether the clastic sediments here were deposited at maximum inundation or at some recessional level. The Bouse deposits are overlain by tributary fanglomerate. At some locations the Bouse grades upward into the tributary gravel, but in other places they are separated by an erosional unconformity. At this location it appears that only a modest amount of erosion of the Bouse Formation occurred upon recession of the large body of standing water, prior to onlapping by tributary alluvial fans. On the south side of Secret Pass Wash we see almost entirely indurated tributary fan deposits, however, with only a small outcrop of Bouse deposits preserved near the base of the cliffs several hundred meters downslope. We infer that the Bouse deposits were substantially eroded there prior to deposition of the thick tributary fan package, but it is not clear whether this fan package correlates with the fan deposits immediately above the Bouse on the north side or is younger. Walk slightly less than 0.3 mi up the wash to where a sizable tributary enters from the north. Here the Bouse Formation and an underlying tephra bed are exposed in a thick package of tributary fanglomerate (Fig. 14). The Bouse Formation consists only of a thin bed of limestone ~10–15 m above the wash. The limestone bed can be traced intermittently from our previous stop to this location and continues on in the narrow canyon to the east. In some exposures upstream, the Bouse Formation consists primarily of massive or cross-bedded sand deposits over a very thin limestone layer, but in contrast to a few hundred meters to the west, the total deposit thickness is 3 m or less. The Bouse deposits here seem to represent a brief incursion of quiet water into an alluvial fan environment, but it is possible that more of the Bouse was removed by erosion prior to renewed tributary fan deposition. Just above the active channel, there is a fairly continuous layer of tephra up to 0.5 m thick and several other large pods of tephra at the same level; these have been identified as the 5.5 Ma Connant Creek tephra via geochemical analysis and correlation. The tephra and Bouse beds are separated by ~10 m of fanglomerate. This site is high up on the margin of the valley, so the 5.5 Ma date provides a maximum constraint for the age of the filling of Mohave Valley with deep water. Directions and Highlights en Route to Stop 1.4 Return west down Secret Pass Wash to Bullhead Parkway. Turn left (S) onto Bullhead Parkway and reset odometer. At ~0.5 mi there are more exposures of the Bullhead unit in roadcuts on both sides of the road. Enter the valley of Silver Creek at

373

Black Mtn gravel QTa Colorado R sand

alluvium of Bullhead City

roundstone gravel Colorado R sand

Figure 13. An example of alternating sand and gravel beds in the alluvium of Bullhead City. The Bullhead beds are unconformably overlain by coarse tributary fan gravel.

Black Mtn fanglomerate Tfb2 Bouse basal limestone

~10 m

Black Mtn fanglomerate Tfb1

Connant Cr. Cr. tephra

Secret Pass Wash Figure 14. Photograph of the north wall of Secret Pass Canyon showing a thin, fairly extensive bed of the 5.5 Ma Connant Creek tephra and the thin, continuous basal limestone bed of the Bouse Formation separated by ~5 m of local fanglomerate. The altitude of the wash bottom at this site is ~415 m (1360 ft) above sea level.

~1.2 mi. At 1.7 mi, turn east (left) on Silver Creek Rd and proceed up the piedmont on dissected Plio-Pleistocene fan remnants for 1.7 mi. Pull off on the left side of the road. Hike to the bottom of Silver Creek (the broad, deep wash north of the road) and proceed ~100 m (328 ft) down the wash along the south side of the valley.

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Stop 1.4: “Plush Toy” Nomlaki Tephra Site At this site (named for some of the interesting trash that litters the slope) there is a middle Pliocene tephra deposit preserved in tributary fan gravels. It provides a minimum constraint on the inception of Colorado River downcutting following maximum aggradation (Figs. 12 and 15). As was discussed in the introduction, two separate middle Pliocene tephra deposits have been discovered in Mohave Valley. The older Lower Nomlaki tephra (3.6–4.2 Ma) has been found in distal tributary fan deposits that underlie the highest levels of the Bullhead unit and, thus, provides a maximum constraint for the culmination of river aggradation. The older Nomlaki tephra is described at Stop 2.5. Along Silver Creek, the younger Nomlaki tephra (3.3 Ma) is incorporated into tributary gravels that rest on an unconformity on top of Bullhead deposits and, thus, provides a minimum constraint for the timing of incision of the river after peak river aggradation. The lower 10 m or so of section exposed here is primarily sand and gravel deposits of the Bullhead unit, with some interfingered tributary gravel deposits. These deposits are truncated by a slight angular unconformity that is overlain by 5–8 m of tributary fan gravel dominated by local lithologies, with minor reworked Colorado River gravel. There are several “pods” of tephra approximately in the middle of the capping tributary gravels, 125–300 m downslope from the dirt track. The highest outcrop of the tephra bed is 395 m (1300 ft) asl, slightly below the highest level of Colorado River aggradation (400 m asl). Farther down Silver Creek, the same bed is found at 350 m (1160 ft) asl, indicating that the river had incised at least 50 m below its maximum level of aggradation by 3.3 Ma. Thus, the Colorado River reached its maximum level of aggradation sometime after 4.2 Ma and possibly as late as 3.6 Ma, and by 3.3 Ma was substantially incised. Directions and Highlights en Route to Stop 1.5 Proceed for ~2.7 mi up Silver Creek Road, passing a low bedrock hill on your left. An outcrop of tufa associated with the

Bouse Formation is preserved on the side of this hill at an altitude of 260 m (1785 ft) asl. This is the highest tufa outcrop associated with the Bouse that has been found (Metzger and Loeltz, 1973), and its altitude is consistent with the Bouse deposits at Stop 1.6. Continue 0.2 mi farther east and pull off on the left side of the road. Follow the dirt road to the bottom of Silver Creek and cross the valley to view outcrops of the Bouse limestone intercalated between alluvial fan deposits. Stop 1.5: Silver Creek Bouse Site The purpose of this stop is to view the highest extensive Bouse deposits found along the lower Colorado River. This is also a classic, enigmatic exposure of a limestone a few meters thick sandwiched between relatively coarse alluvial fan deposits. The limestone is draped over cobbles and boulders that mantled an alluvial fan surface (Fig. 16). In the westernmost part of the outcrop, the Bouse is fairly pure, fine-bedded white limestone. There is a clastic component interbedded with limestone farther to the east and higher up in the section, and at the easternmost exposure there is abundant fine gravel. The altitude of the base of the Bouse ranges from ~530 m (1740 ft) asl in the west to 550 m (1810 ft) asl in the east. The contact between the limestone and the overlying fanglomerate is clearly erosional in some places; for example, notice where the fan gravels fill small channels cut into the Bouse deposits. In other places fine gravel beds and limestone are intercalated and the transition appears to have been more gradual. Substantial fan aggradation continued after the interval of Bouse deposition, as evidenced by the ridge capped by alluvial fan deposits that rises to 585 m (1920 ft) asl northeast of this site. The tranquil water into which the Bouse Formation was deposited here surely represented a dramatic and (likely) brief departure from the subaerial alluvial fan conditions that had previously dominated the upper piedmont. It is interesting to note here how little-disturbed the underlying alluvial fan surface appears to have been when it was inundated. If the Bouse Formation represented a marine incursion into this area, one might

Qtrash QTa

Ttn

Tbh

Silver Creek valley bottom

Figure 15. Photograph of a bed of the 3.3 Ma Nomlaki tephra (Ttn) in tributary gravel deposits at the “Plush toy” site. The tributary deposits rest above a slight angular unconformity on the sand and roundstone gravel of the alluvium of Bullhead City. Tbh—alluvium of Bullhead City (Colorado River deposits); QTa—locally derived alluvial fan deposits derived from the Black Mountains.

Birth of the lower Colorado River expect to see a more elaborate transgressional sequence including beach deposits overlain by quiet water deposits as water depth gradually increased. Even wave action along the shore of a gradually rising lake would probably disturb the underlying fan surface, although this location might have been sheltered by the bedrock hill to the west. It is possible that the maximum level of inundation in the valley was short-lived and possibly modulated by short-term variations in regional climate. Day 2: Start and End in Laughlin Directions and Highlights en Route to Stop 2.1 Travel north on S. Casino Drive through Laughlin and turn left (W) onto NV 163. Prepare to take an abrupt right turn onto N. Casino Drive, follow its sharp curve back to the east and then continue north toward Davis Dam. At ~0.6 mi park along the side of the road (see Fig. 5). Hike due west through the abandoned gravel pit. This pit served as an aggregate source during construction of the dam and its facilities. A large construction camp with many structures once occupied the flat area immediately north of here. Presently, the area serves as home to a small population of transient residents. Please be respectful of their home and their privacy, but also lock your vehicle.

375

Introduction There are extensive exposures of Pleistocene alluvium of the Colorado River near Davis Dam, just north of Laughlin. Our mapping in this area has revealed a series of flood deposits, unconformities, and stratigraphic ambiguities that make unraveling the record and characterizing the Chemehuevi beds a challenge (Fig. 17). Stop 2.1: The Davis Dam Bluffs—Best Dam Stratigraphy Ever This stop highlights various aspects of the Quaternary record of the Lower Colorado River including evidence for a postintegration catastrophic flood through the Pyramid hill area and important discontinuities in distinctive packages of Pleistocene Colorado River deposits. Davis Dam was constructed between 1942 and 1945 and 1947–1950 (the brief hiatus owing to material shortages during World War II). It is an earth-fill dam. Its base sits on river alluvium at a depth of 45 ft below the streambed. Exploratory borings made during dam planning and construction indicate that the depth to bedrock at the dam site is 200 ft (Bahmoier, 1950;

0

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relict boulders and cobbles on paleosurface Black Mtn fanglomerate Tbf1

Qcg

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Figure 16. Photograph of the basal Bouse limestone intercalated between indurated tributary fanglomerates on the north bank of Silver Creek. The lower Bouse beds are draped over cobbles and small boulders on a relict alluvial fan surface. At this location, there are many thin sandstone and siltstone beds in the upper part of the Bouse exposure. The upper contact of the Bouse Formation here is mildly erosional, but in other parts of this exposure the upper fanglomerate fills small channels cut into the Bouse beds. Tbf1—pre-Bouse fanglomerate of Black Mountains; Tbf2—post-Bouse fanglomerate of Black Mountain.

Qcl2

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basal

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Qx

Figure 17. Detailed geologic map of the Davis Dam bluff vicinity (modified from Faulds et al., 2004); Ydg—Davis Dam granite; Tri—Rhyolite intrusion; QTg—undifferentiated Plio-Pleistocene river gravel; QTa— Plio-Pleistocene alluvial fan gravel; QTr—Riverside beds; QTlc— Laughlin conglomerate; Qao—early Quaternary alluvial fan gravel; Qc—Chemehuevi beds (Qcg—river gravel; Qcl—lower Chemehuevi beds; Qcl1—mud unit 1; Qcl2—mud unit two; Qcu—upper Chemehuevi beds; Qcm—Mohave sediments); Qai—intermediate age alluvial fans (middle to late Pleistocene); Qr1—alluvium of Riviera (Holocene).

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topography and a paleosol in the Laughlin conglomerate indicate prolonged subaerial exposure of the unit prior to burial by a series of disconformable fluvial units. We interpret the Laughlin unit as the result of a large flood that eroded a weak bedrock outcrop in the Pyramid hills area adjacent to the main channel. Exposures of the deposit define a lobe that widens downstream from a paleochannel in the large outcrop of Proterozoic granite southwest of Davis Dam. The Laughlin conglomerate has a total thickness of ~20 m (70 ft), and we have found no evidence to suggest that it is comprised of more than one deposit. Erosional topography on the Laughlin unit and underlying deposits is overlain by the base of the Chemehuevi beds, which form the core of the two major fluvial terraces at this site. The thickest part of the sequence backfills the lower end of the paleochannel that was the conduit and likely source of the Laughlin conglomerate. The higher terrace has a maximum elevation of 252 m (830 ft), but the maximum traceable extent of the underlying deposits reaches to 256 m (840 ft) on the Newberry piedmont. The lower terrace surface evident here is predominantly a strath. It has a maximum elevation of 225 m (740 ft). The riser separating the two terraces is mantled with colluvium, and both terrace surfaces have locally extensive mantles of eolian sand. Proceed from the overview spot toward the steep bluffs. The central part of the bluffs is covered with loose, sandy veneer. Just beyond the covered section to the south are a series of alluvial fan remnants protruding from the base of the bluffs. Hike up the

U.S. Bureau of Reclamation, 1955). The dam is named in honor of Arthur Powell Davis, director of the Bureau of Reclamation from 1914 to 1932, and an important figure in the development of the Colorado River’s water resources. Hike west from N. Casino Drive for a few hundred yards and climb onto one of several terrace remnants for a good, panoramic view of the deposits that form the core of the bluffs. (Fig. 18). The base of the section includes sparsely exposed, indurated remnants of Pliocene or early Pleistocene fanglomerate. Late Miocene Newberry Mountains–derived fanglomerates are exposed immediately east of the river. The basal fanglomerate unit is overlain by the Riverside beds, a Colorado River deposit of mud, sand, and gravel. Thinly bedded, flat-lying fine-grained components of the Riverside beds are exposed along N. Casino Drive and are similar in composition and appearance to overlying river deposits. Beds of fluvial gravel and gravelly sand in the Riverside unit are exposed locally. The Riverside beds are unconformably overlain by the conglomerate of Laughlin (informally called the Laughlin conglomerate), a coarse conglomerate comprised of cobbles and boulders of locally derived Proterozoic granite mixed with well-rounded, far-traveled pebbles and cobbles. The Laughlin unit is texturally similar to the transitional Pyramid gravel, but with the important distinction that the Pyramid unit does not contain exotic clasts. The Laughlin conglomerate was deposited well after the river had incised through the Bullhead fill to near the pre-integration valley axis. Otherwise, its age is not well constrained. Erosional

Soil

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Figure 18. Annotated photo mosaic of the Davis Dam bluffs, north of Laughlin, Nevada. Qao—older (early Pleistocene?) alluvial fan deposits; Qao/Qcl—interfingered fan gravel and Colorado River sediment; Qcls—immature fine sand of lower Chemehuevi beds; Qcl1—mud unit 1 of the lower Chemehuevi beds; Qcl2—mud unit 2 of the lower Chemehuevi beds; Qcu—medium sand and gravelly medium sand of the upper Chemehuevi beds; Qe—veneer of eolian sediment.

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Directions and Highlights en Route to Stop 2.2 Upon leaving Laughlin, proceed west about a mile to Needles Hwy, and turn left (S) ~22 mi along the east flank of Mohave Valley to Needles, California, where you will enter I-40 eastbound (toward Kingman,. Arizona). Proceed southeast on I40 seven mi through and past Needles (Figs. 4 and 19). On the right is the domal Sacramento Mountains metamorphic core complex, and ahead is the Chemehuevi Mountains

core complex, two elements of the Colorado River extensional corridor. In between, at ~two o’clock, a lone peak exposes the Miocene Sacram plutonic suite of Campbell and John (1996) which was intruded during the extension. Mafic plutonic rocks of Miocene age in this suite are part of an elongate zone over 200 km long of high residual isostatic gravity highs that track the axis of the Colorado River extensional corridor (Simpson et al., 1990). Carr (1991) speculated that isostatic subsidence of the dense rocks that cause the gravity high may have influenced the Colorado River’s southward course. Beyond the agricultural inspection station, part of I-40 is built on swelling clays of the Bouse Formation. Swelling of the clay after heavy rains in the early 1990s made the highway unsafe and required expensive repairs. A thin bed of white Bouse marl is present in this area, and extensive exposures can be seen ahead of horizontal-bedded mud and sand of the Bouse. When you have been on I-40 for ~6 mi, the route passes near the Park Moabi exit and the river is visible. Recent drilling efforts here related to groundwater remediation efforts have provided new information on the subsurface stratigraphy, including identification of tens of meters of river laid sand recovered from below the west bank of the Colorado River. The upper part of this sand section is believed to be historical, as the river here aggraded ~8 m in the mid–twentieth century in the backwaters of Lake Havasu (Metzger and Loeltz, 1973). A wood fragment collected from fluvial sand at a total depth of 18 m (60 ft) recently yielded a mid-Holocene 14C age. This supports the conclusions of Metzger and Loeltz (1973) and Metzger et al. (1973) that the predam river aggraded ~20–30 m in the Holocene. They based that conclusion on Holocene-age wood fragments recovered from drilling in fine-grained fluvial sediments beneath the floodplain 100 km downstream, and from a similar section drilled 15 km north of this spot.

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spine of the northernmost remnant and peer into the precipitous gully (just left of center in Fig. 18). The base of this section is a fining-upward sequence of alluvial fan gravel from the Newberry Mountains (Qao) and interfingered sand of possible Colorado River origin. This package grades upwards into a massive bed of flat-lying, immature fine sand that comprises the base of the lower Chemehuevi beds. The composition of these sands suggests a largely local derivation, but their texture and sorting suggest reworking by the Colorado River. The immature sand is overlain by flat-lying, slightly red fluvial mud which is, in turn, unconformably overlain by a very similar looking deposit of mud. This unconformity within the lower Chemehuevi beds is obvious here, but it becomes cryptic toward the south where it is flat and unremarkable in many places. Here, the mud-on-mud contact is marked by a sandy, homogeneous slab-like feature, likely a soil formed in finegrained hillslope colluvium. The irregular trend of the “slab” suggests a furrowed, badland-type erosional setting. The mudon-mud unconformity represents a hiatus in river aggradation followed by subaerial exposure and erosion which was followed by a second episode of aggradation of fluvial mud. The two mud units are unconformably overlain by a thick sequence of clean, medium fluvial sand and minor gravel (the upper Chemehuevi beds). This contact can be traced from the north for several hundred meters to this point. Just beyond this point, it abruptly drops ~10 m, and then continues as a flat contact for several hundred meters to the south. The steep drop in the sand-on-mud contact suggests a channel margin. The continuity of stratigraphy in the upper unit indicates a thick net aggradation event. We have identified a paleosol in the lower part of the upper sand unit that indicates some complexity in the deposit’s history. The alluvial fan unit that is graded to the top of the Davis Dam terrace (Qai1) is only clearly coeval with the gravelly sand component of the upper Chemehuevi beds. Relatively weak soils in the fan surface (distinct Bw and stage II to ~stage III Bk horizons) suggest that the upper Chemehuevi beds are late Pleistocene. Similar soils were dated at ca. 60 ka in the Parker area (Ku et al., 1979). The ages of the underlying mud beds are, however, not fully resolved, and we suspect that the lower Chemehuevi mud unit here is much older than the upper sand unit, while the younger mud unit may be closer in age and may form the base of a floodplain package that flanks the lateral margin of the thick sandy unit. Luminescence ages from a similar sequence of muds near Cottonwood Landing (34 km north) range between 40 and 70 ka (Lundstrom et al., 2000, 2004).

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After 6.5 mi, cross the Colorado River on I-40 into Arizona. You will see the Needles on the right, pinnacles of synextensional, steeply west-tilted lower Miocene volcanic and sedimentary rocks (Howard and John, 1997). To the left (NE) of the Needles, the basement substrate of this section is exposed to paleodepths of 8 km below the Miocene rocks. All of these rocks were tilted and structurally superposed by eastward fault slip on the Chemehuevi Mountains metamorphic core complex in Miocene time (John, 1987; Miller and John, 1999). Gently east-sloping rock surfaces on the California side of the river indicate the position of the exhumed bounding Chemehuevi detachment fault, which projects eastward under the allochthonous rocks of the Needles. The river’s path follows near this structural boundary as it courses southward through Topock Gorge. Continue 0.8 mi past the river to Exit 1, exit right and turn left over the freeway toward Topock. Proceed 4.4 mi. Much of the floodplain of the Colorado River in this area is now occupied by Topock Marsh, which flooded as a consequence of the river aggradation in response to Parker Dam downstream. From the fork at Golden Shores, keep left and proceed 2.7 mi. Turn right on an obscure jeep path, and proceed up the wash 0.2 mi to an area traversed by buried pipelines. Stop 2.2: Piedmont Fans Interrupt the Colorado River’s Fun From here, take a walking traverse ~3 km round trip up the main wash to the east and a tributary to the northeast. Along the transect, one can examine well-exposed stratigraphy, compare older and younger Colorado River deposits and intervening fanglomerate, and consider depositional environments and geomorphic response to at least two successive episodes of river incision and aggradation. The stratigraphy here matches that described by Metzger et al. (1973) and Metzger and Loeltz (1973). The high banks of the main wash expose >20 m of cross-bedded fluvial sandstone and interbeds of fluvial roundstone conglomerate. These deposits can be traced to exposures a mile and a half to the north-northeast that were identified by Metzger and Loeltz (1973, their Figure 14) as their unit B, which we consider equivalent to the Pliocene-age alluvium of Bullhead City of this report. Look for well-rounded pebbles of chert, quartzite, and Paleozoic limestone derived from far upstream. Locally derived angular or subangular clasts of volcanic rocks and gneiss make up at least half of the larger clasts. Well rounded quartz sand grains in the sandstone may be reworked from Permian and Jurassic sandstones on the Colorado Plateau. Rusty zones in the deposits here are commonly associated with clay balls, wood casts, or vertebrate remains. The clay balls may derive from bank erosion of rare clay beds in the sequence, or from the Bouse Formation. Note the planar crossbeds in both sandstone and conglomerate. Do the thickness of crossbeds indicate flow depth and large discharge of a braided stream? One km to the southwest, the fluvial sand overlies basaltclast fanglomerate that may either predate or postdate the Bouse Formation. It may correlate with a fanglomerate 100 km to the south that Metzger et al. (1973) recognized (their unit A) as over the Bouse Formation and unconformably overlain by unit B alluvium (our Bullhead City unit).

Note that the cross-bedded fluvial sandstone is exposed in the bed of the modern wash. This indicates that the modern flatfloored drainage is graded to Topock Marsh on a pedimented surface, and by inference this particular wash postdates the 20–30 m of Holocene aggradation known for the nearby Colorado River floodplain downstream. Capping the alluvium of Bullhead City is a paleosol overlain by dark, coarse-grained alluvial-fan deposits (fanglomerates) dominated by basalt derived from the Black Mountains. Metzger and Loeltz (1973) described a series of such fanglomerates (their unit C, piedmont gravels) as alluvial fans that prograded into the valley as the thick underlying fluvial aggradation package (unit B, our alluvium of Bullhead City) underwent incision. Several ages and terrace levels of such fans are preserved in the landscape east of us. The Bullhead City unit into which these fans are inset can be traced intermittently to the east from here to valley-flank elevations as high as 200 m above the river, indicating the thick valley fill the unit represents. Walk up a narrow tributary to the NW and see that the Bullhead City unit here, slightly more distal from the valley axis, lacks the two conspicuous conglomerate layers that are nearer to the parking area, and instead exposes mainly sandstone and 3–4 clay lenses as thick as 0.3 m. Do these clays record standing ponds in distal parts of a braidplain? Sandstone regionally dominates the Bullhead City unit. Resistant roundstone pebbles derived from less thick interbedded conglomerate layers form conspicuous lags and are commonly reworked into alluvial fans and soils. At the head of the gully, pale orange layered mud and very fine sand (unit D of Metzger and Loeltz, 1973) of the Chemehuevi beds overlies the sandstone, locally with a thin intervening alluvial-fan deposit of basalt boulders (unit C) or a pebbly paleosol. One km to the NW, 2 m of calcite-cemented paleosol lies at the top of the fanglomerate, indicating long exposure before the Chemehuevi Formation was deposited. The overlying Chemehuevi layered mud unit is many meters thick at this position, but thins to 0.5 m within 1.5 km to the north. Does the mud represent overbank deposits on a floodplain? Here, as elsewhere, the mud is overlain by well sorted, unconsolidated, light-toned sand, which forms gentle slopes. The sand, unit E of Metzger et al. (1973) and Metzger and Loeltz (1973), is ~20 m thick and locally exhibits two internal red pebbly paleosols. A lag of roundstone and angular pebbles caps the sand. Nearby (1 km from here) a thin (0.5 m) basalt-clast fanglomerate intervenes between the loose sand and thin underlying mud of unit D. This stratigraphy suggests that an alluvial fan prograded into the valley as the mud deposition ceased. Climb to the top of the sand for an overview of the stratigraphy and of the preserved fan morphology upslope of the fanglomerate deposits that interfinger in the section. A succession of unit C fans, graded to increasingly lower topographic levels, may correspond to intermittent incisional lowering of the valley floor after unit B aggraded, or to smaller episodes of intermittent river aggradation during or following the major lowering. Muds and sands of the Chemehuevi beds (units D and E) represent one

Birth of the lower Colorado River or more cycles of aggradation and subsequent degradation of the river valley. The best existing dates for these beds are late Pleistocene, between 35 and ca. 100 ka (Bell et al., 1978; Blair, 1996; Lundstrom et al., 2004). Directions en Route to Stop 2.3 Return to vehicles and to Hwy 95. Turn left and backtrack 2.7 mi, passing back through Golden Shores and the fork for the Oatman Hwy. Turn left on to Polaris Road, and proceed east for 1.5 mi (Fig. 19). Park at the crest of a small hill. Stop 2.3: Mammoth Retires at Golden Shores At this stop, the stratigraphic setting and implications of an early Pleistocene mammoth site can be examined and considered. Mammoth remains were first reported from the Colorado River valley by Newberry (in Ives, 1861). He found a tooth at what the expedition called Elephant Hill, now in Lake Mohave. Other sites up and down the river valley have been discovered over the years, and the total now exceeds 12 sites from the Lake Mead area to Mexico (Agenbroad et al., 1992). Many of the specimens have been designated as Mammuthus columbi, whereas others have been designated as the older Mammuthus meridionalis. A nearly complete skeleton was recovered at this site. The mammoth evidently settled on its back in shallow water and was encased in fine clays; abundant imprints of reeds and sedges in the clay suggest a marshy environment (Agenbroad et al., 1992). Unfortunately, teeth were removed from the remains and not available for examination. Agenbroad (L.D. Agenbroad, 1995, 2005, personal commun.) evaluated a photograph of the teeth and concluded that it was probably a M. meridionalis, and dates to ca. 1.5–1.7 Ma. The Golden Shores mammoth site occurs at an elevation of 207 m (680 ft) asl within a section of interbedded and interfingering Colorado River gravel, sandstone, mudstone, and locally derived angular gravel containing some reworked roundstone pebbles. Nearby within the deposits, root casts below a minor paleosol indicate that deposition of the fluvial section was discontinuous. The mammoth site lies south of the road across a gully, and is marked by a post and a sign that warns against disturbance. Inasmuch as North American mammoths are not known before the Pleistocene (L.D. Agenbroad, 1995, personal commun.), the mammoth age and the facies assemblage suggest deposition during an early Pleistocene aggradational episode of the river. If so, this postdates the alluvium of Bullhead City and predates or may overlap in age with some of the Chemehuevi beds or the older, post-Bullhead riverside beds. Alluvial-fan deposits that lie below the mammoth horizon contain roundstone pebbles that suggest derivation from older (Bullhead City) river deposits. From the mammoth site to 5 km to the east are discontinuous exposures of fluvial sand and gravel, which we regard as mostly B, the alluvium of Bullhead City. There are many such exposures of unit B in this part of the valley. The exposures east of the mammoth site extend to a hilltop 5 km east where reworked (unit B) roundstones in a soil reach an eleva-

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tion of 311 m (1020 ft) asl, over 100 m higher than the mammoth site. Assuming these are older Bullhead deposits, they provide a potential source for later reworking of roundstones into alluvialfan deposits that interfinger with or underlie the fluvial section at lower elevation that hosts the mammoth. Do the deposits that hold the mammoth represent a local fluvial aggradation inset into the older, thick Bullhead City sequence following a major late Pliocene or earliest Pleistocene incision? If so, this hints that, in addition to the aggradation recorded by the Chemehuevi beds, the river experienced at least one gravel-rich early Pleistocene aggradational episode. Directions and Highlights en Route to Stop 2.4 Turn around, return south 1.5 mi, and turn right on Route 66. In 0.7 mi, fork right (straight) on the Oatman Hwy (Fig. 19). This road is part of historic Route 66, famed in song and immortalized in John Steinbeck’s The Grapes of Wrath. As you pass through the community of Golden Shores, ponder the desperate people fleeing the 1930s dust bowl en route to California. Two and a half miles farther, the road crosses extensive deposits of red mud of the lower Chemehuevi beds. Then, 1.7 mi farther, just past a major gully, park on the right side of the road at a bend to the left. Stop 2.4: The Crinkled Quaternary Alphabet The object of this stop is to discuss evidence for deformation in Mohave Valley since the inception of the Colorado River. Compared to southwestern California, deformation features younger than Miocene are uncommon along the lower Colorado River valley. This relative tectonic quiescence made the region attractive in the 1970s for site-proposals for nuclear power plants downstream near Vidal and Blythe. The plants were never built, but related geologic investigations added much to the knowledge of late Cenozoic geology in the region. The Golden Shores area of Mohave Valley is one of the rare areas where several faults and folds deform post-Bouse deposits. Tilting and faulting were recognized by Metzger and Loeltz (1973) 9 km south of this stop in their unit B (Bullhead alluvium of this report), at a locality earlier photographed by Lee (1908). Based on the presence of faults and tilted beds, and of gravels logged from a well 95 m (310 ft) below the floodplain of central Mohave Valley, Metzger and Loeltz (1973) suggested that their unit B (our Bullhead City unit) may be structurally sagged beneath Mohave Valley. Pleistocene alluvial fan deposits are clearly faulted in the Needles graben a few miles southeast of this site (Fig. 20; Purcell and Miller, 1980; Pearthree et al., 1983). Looking east across the gully from Stop 2.4, view an exposure of sandy-gravel alluvial-fan deposits that dip southwest 20° and are cut by a pair of conjugate WNW-striking normal faults. The deformed deposits are capped unconformably by a basaltclast fanglomerate, which exhibits >2 m thick calcic soil horizon that is not obviously deformed. This suggests that the deformation age was earlier than middle Pleistocene. The undeformed capping fanglomerate forms a partly preserved fan terrace (C2 on Fig. 20) that is one of a downward

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Figure 20. Deformation features near Golden Shores, Arizona. Air photo of (1) Needles graben, (2) monocline (double arrow symbol), and (3) small thrust fault (with teeth) northeast of Golden Shores, Arizona. These features deform an old fan (C3) that is one of several that postdates the alluvium of Bullhead City. The C3 fan is inset into the topographically higher and older C2 fan, which unconformably overlies dipping beds (dip symbols) indicating older (Pliocene?) deformation. Small offsets along the Needles graben trace southeastward into younger alluvium (upper Quaternary), indicating that deformation on the graben faults continued into late Pleistocene or Holocene time.

succession of post-Bullhead fans from the Black Mountains graded toward the central part of the valley. Metzger and Loeltz (1973) also reported this observation in relation to their stratigraphy. A mile ESE of this location, the Needles graben and two other structures deform at least 2 younger Pleistocene alluvial surfaces (Fig. 20). The most prominent faulted surface is also capped by thick calcic soil and exhibits several deformation features: (1) the NW-striking Needles graben (Purcell and Miller, 1980; Pearthree et al., 1983), (2) a NW-striking ramp interpreted here as a monocline, and (3) a thrust fault that strikes perpendicular to those features. There is evidence for recurrent fault movement during the middle and late Quaternary along the graben (Pearthree et al., 1983), and the large scarp associated with the monocline (~25 m) is consistent with recurrent deformation. Within a few kilometers of here are also several other faults and a syncline that deform the alluvium of Bullhead City or younger deposits. Quaternary graben structures have also been reported in alluvial fans near the river 100 km south of here near Blythe and 30 km south in Chemehuevi Valley (Purcell and Miller, 1980). The cause of the local extension recorded by these grabens is uncertain. The Needles graben and related monocline and small thrust directly overlie a gravity low indicative of thick low-density sedimentary fill (Gray et al., 1990). Sedimentary compaction of this fill may explain the graben and monocline. The shape of the gravity low, however, does not offer an obvious explanation for the bulk down-

throw to southwest, or for possible sagging of sediments under the river valley as suggested by Metzger and Loeltz (1973). Highlights and Directions en Route to Stop 2.5 Continue 6 mi NE on Hwy 66. Turn west (left) onto dirt track and head nearly due west down the piedmont (Fig. 19). The fine-grained sediment that is extensively exposed on the north side of this valley may be the Bouse Formation, although no basal limestone is evident here. The altitude of these deposits is between 378 and 390 m (1240–1280 ft) asl, so they could be related to the Bullhead unit, but no cross-bedded sand or roundstone was observed in this outcrop. Continue west for 1.4 mi and turn north (right) onto the dirt track that follows the powerline. The road traverses a moderately to darkly varnished late (?) Pleistocene alluvial fan deposits, with limited Holocene deposits along active washes. Much higher ridges east of the powerline are probably capped with Plio-Pleistocene fan remnants. The first limited exposure of the Lower Nomlaki tephra is visible about 1.8 mi along the powerline on the eastern nose of a ridge remnant west of the dirt track. Continue north ~2.5 mi and stop on this low ridge beneath the powerline. Stop 2.5: Fantastic Tephra Outcrop Several years ago, we made a fortuitous discovery of a thin tephra bed interbedded with fan and river deposits near the 13th

Birth of the lower Colorado River green of a planned golf course near Bullhead City. This discovery helped us reframe our conception of the river’s early history. We subsequently found (or in one case, rediscovered) the same tephra in two additional locations, along the powerline road north of Topock and northwest of Cottonwood Landing, Nevada, in Cottonwood Valley. All of the tephra outcrops are within 30 m of the highest extant gravels of the Bullhead alluvium. Geochemical fingerprinting of tephra samples from each site indicates that they are the Lower Nomlaki tephra, which is dated between 3.6 and 4.2 Ma (A. Sarna-Wojcicki, 2002, personal commun.). The Nomlaki series tephra beds are members of the Tehama and Tuscan Formations of the Sonoma Volcanic Field in north central California (sources cited in Sarna-Wojcicki et al., 1991). Because of its presence high in the Bullhead aggradation sequence, this tephra provides an important constraint on the timing of maximum Colorado River aggradation in this valley. The area immediately around us contains extensive exposures of the Lower Nomlaki bed. A tephra exposure in this area was first noted by Metzger, who described it to J. Bell in the early 1970s. At that time its chronologic and stratigraphic importance was not fully appreciated. Tephra is intermittently exposed on the sides of at least 3 parallel ridges for ~1.5 km in a north-south direction. Altitudes of the exposures range from ~365 m (1200 ft) to 370 m (1215 ft) asl. Everywhere we have observed the tephra bed in this area it rests on sand, silt, or fine gravel and it is overlain by similar deposits; no definitive Colorado River deposits have been identified immediately above or below the tephra outcrops here, but limited exposures of Colorado River gravel and sand deposits are on this ridge that lie stratigraphically above the tephra and there are outcrops of clean sand and river gravel that appear to underlie the tephra bed ~0.5 km west of the powerline road. Because the tephra bed is so extensive here, we can reconstruct the shape of the surface on which it was deposited. This paleosurface dipped gently (~0.5°) to the west or southwest. Based on the character of the deposits surrounding the tephra and the surface slope, we infer that the tephra was deposited on a gently-dipping, relatively fine-grained distal alluvial fan not far from the Colorado River floodplain. Subsequent to tephra deposition, the Colorado River continued to aggrade and eventually onlapped the distal alluvial fans, depositing limited river sand and gravel at least as high as 385 m (1260 ft) asl in this area. The geologic and geomorphic settings of the other two exposures of Lower Nomlaki tephra were also apparently marginal to the Colorado River floodplain. In each case, the tephra is found ~400 m (1320 ft) asl. In the 13th green exposure, the tephra bed rests on fan gravel but is overlain by gravel that includes roundstone gravels. At the Cottonwood Landing site, the tephra bed is at the base of a tabular bed of Colorado River sand interfingered with fan deposits. The tephrochronologic constraints on the fluvial transition in Mohave and Cottonwood valleys (as defined by the pre–Bouse Connant Creek tephra and the Bullhead peak aggradation Lower Nomlaki tephra) nearly perfectly mirrors what is currently

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known about the integration and incision of the Colorado River through the Grand Canyon. The same Connant Creek bed is beneath river-precluding lacustrine deposits in the western Lake Mead area (Faulds et al., 2002). At Sandy Point, Arizona, near the mouth of the Grand Canyon, a recently redated 4.4 Ma basalt flow is intercalated with river gravels 100 m above present river grade, so the Colorado River was obviously well-established and probably aggrading west of Grand Canyon at 4.4 Ma (Howard et al., 2000). These new geochronologic constraints support the process-linkage between canyon incision upstream and massive valley aggradation downstream. Casting the mechanism for the deposition of the alluvium of Bullhead City in this new light is an important step in reframing the problem of the evolution of the river. It appears that the process of drainage integration forced a huge pulse of fluvial aggradation, and Cottonwood and Mohave valleys were buried, not carved by the Pliocene Colorado River. The deposition of the Bullhead unit was a unique event in the history of the river that was driven by an internal adjustment to a particularly voluminous, relatively coarse sediment load. Thus, the character and volume of sediment deposited by the river in the early Pliocene is quite different from any depositional interval during the Quaternary, when aggradation was presumably driven by external controls such as climate change. Day 3: Laughlin to Las Vegas Via Detrital Valley and Hoover Dam Introduction Driving distance ~180 mi. Stops during the first half of Day 3 are in Cottonwood Valley and contain evidence for the body of standing water that formed above and eventually breached the Pyramid hills divide, which filled both Mohave and Cottonwood valleys with a deeper and larger body of standing water. Stops during the second half of the day involve looking at distant outcrops of the late Miocene Hualapai Limestone and high-level Pliocene Colorado River deposits in Detrital Valley, and a spectacular set of high-standing potholes carved in bedrock near Hoover Dam. Directions and Highlights en Route to Stop 3.1 Depart Laughlin by heading north on Casino Drive and turning east at the intersection with Hwy 163. Bear left (N) at the intersection of the Bullhead Parkway and AZ 68 (0 mi) (Figs. 1 and 4). At 11.4 mi the route crosses Union Pass into Sacramento Valley. Continue 25.5 mi to the Junction with U.S. 93 and take the exit for Boulder City/Las Vegas, Nevada. At ~39 mi, the route crosses the low divide between Sacramento and Detrital valleys near the turnoff to Chloride, Arizona. Exit left at the junction with Cottonwood Road at 46 mi. Follow Cottonwood Road for 8 mi through a sparse residential area to the Lost Cabin Spring turnoff to the left (not marked). Follow this dirt track (high-clearance best but not absolutely necessary) for 3.2 mi to Lost Cabin Spring and continue south for 5.8 mi until the road forks at upper Lost Cabin Wash. Take the right fork down the wash (W). The

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first stop is the base of a prominent west facing bluff of indurated fanglomerate ~5.4 mi down the wash. Stop 3.1a: Miocene Fanglomerate. Many of the deposits exposed in Lost Cabin Wash evidently predate any influence from the developing Colorado River. The lowest part of the section consists of relatively fine alluvial-fan deposits composed almost entirely of clasts derived from the Newberry Mountains to the west, indicating that the preriver era fans extended from the mountain front to ~3 km east of the Arizona shore of Lake Mohave. Locally these fan deposits overlie older, tilted fanglomerates. The beheading of the Newberry Mountains–derived fans by the early river provided copious amounts of local gravel for fluvial transport and redeposition as the lower part of the alluvium of Bullhead City. There are exposures of reworked and intricately cross-stratified deposits of Newberry sand and gravel at the base of the Bullhead unit in southern Cottonwood Valley. Moving upsection to the east, we see a gravelly fluvial sequence that is dominated by Black Mountains volcanic clasts but also contains clasts from the Newberry Mountains. Structures in the gravels and their bedding suggest a NS alignment of the axial valley system similar to the modern one (minus the Colorado River). Stop 3.1b: The Lost Cabin Bluffs. Continuing up the wash, the axial-valley gravel transitions into a prominent, bluff-forming sequence of interbedded sandstone and mudstone that we (informally) call the Lost Cabin beds (Fig. 21). They are a key piece of the latest Miocene stratigraphic puzzle in this area (Fig. 2C). The beds consist of ~100 m of fine-grained, flat-bedded clastic deposits. In locations near here to the north, the 5.5 Ma Connant Creek ash bed occurs within the upper third of the sequence (Fig. 22). This is the same tephra bed viewed at Stop 1.3. The Lost Cabin beds indicate a change in depositional conditions in southern Cottonwood Valley that preceded deep inundation and deposition of the Bouse Formation. More importantly, similar conditions did not exist in northern Mohave Valley, on the other side of the Pyramid hills. We interpret these relations as a strong case for the

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formation of a pre-Bouse body of standing water in Cottonwood Valley, eventual catastrophic drainage through the Pyramid hills, and inundation of both valleys by a second, deeper body of standing water into which the Bouse Formation was deposited. Continuing upsection while driving up the wash, look carefully along the south wall and you can see deposits of mud (Bouse?) in small erosional niches at the top of the Lost Cabin beds. Also evident are some intervening alluvial fan deposits derived from the Black Mountains, and then some unequivocal outcrops of the Bouse Formation appear as prominent tabular beds of buff to cream-colored sandstone and bright white lenses of marl, located mainly along the south wall of the wash. The fine-grained Bouse deposits are intercalated with tributary gravel deposits and possible turbidite deposits. It appears that this area was the focus of a dynamic interaction between the margins of the water body and the tributary alluvial fans. Stop 3.1c: Thin, tenacious outcrops of Bouse marl. More exposures of the Bouse Formation are visible as you continue up the wash. At ~3 mi upstream of Stop 3.1a, a small tributary enters the main wash from the south. Park your vehicle and walk up this wash for ~0.2 mi. In the banks of this wash, the Bouse marl occurs as a very thin bed draped over a likely wave-worked paleosurface. The bed is generally less than 5 cm thick, but it can be easily traced up the tributary. We have noted fossil plant impressions in some of the marl, but have not investigated them in detail. Stop 3.1d: Lost Cabin Beach. The final stop in Lost Cabin Wash is an enigmatic outcrop of cross-stratified, clean sand hemmed in between a bedrock outcrop and indurated, tilted Miocene fanglomerates along the margin of the wash. This site is ~5 mi from the Stop 3.1a. Beds of moderately well-sorted, locally derived gravel and some poorly exposed beds of greenishyellow mud are also associated with the sand. We think that these sands and gravels are potential Bouse shoreline deposits. If this is the case, the fact that they are found at an elevation consistent with the highest Bouse deposits in Mohave Valley indicates a flat water surface at 550–560 m (1800–1840 ft) asl over an axial distance of more than 30 km, thus bearing no evidence of significant regional tilting. These are the highest Bouse-like deposits that we have yet found in Cottonwood Valley. About 2 km west of here we have found a series of unequivocal Bouse sandstone outcrops comprising a fan-delta sequence at ~490 m (1600 ft) asl. Just down the wash from this site there is a distinct, nearly flat-lying shelf that crops out along the base of the fanglomerate bluffs along the south wall. We suspect that this feature may also be related to shoreline processes.

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Figure 21. Oblique photograph of the Lost Cabin bluffs in Lost Cabin Wash, Arizona. Tfb—Black Mountains–derived fanglomerate; Tbl— Bouse limestone-marl; Tlc—Lost Cabin beds.

Directions and Highlights en Route to Stop 3.2 Retrace the route to U.S. Hwy 93 in Detrital Valley, and turn left (N) toward Las Vegas (Figs. 4 and 23). Note how high the alluvial fill of Detrital Valley is perched above the adjacent deeply incised Colorado River valley to the west. Was the Colorado valley equally filled following Miocene extension and subsequently exhumed? Proceed 22.8 mi to the Temple Bar Road and turn right toward Temple Bar. Proceed north 5 mi and park.

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Figure 22. Photograph of the 5.5 Ma ash layer in the upper part of the Lost Cabin beds near Lost Cabin Wash, Arizona.

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Stop 3.2: Pliocene Fluvial Detritus in Detrital Valley This stop (Fig. 24) presents an overview of sediments and landforms involved in the early history of the river in the Lake Mead region. Detrital Valley debouches into Lake Mead (formerly the Colorado River), visible ahead to the north. The valley is one of several north-striking structural valleys that the river historically traversed westward across the Basin and Range province from the edge of the Colorado Plateau, before turning south in the area now occupied by Hoover Dam. The earliest history of the Colorado River in the Lake Mead area and the mouth of the Grand Canyon to the east revolves around the upper Miocene Hualapai Limestone, the youngest unit to predate the river’s arrival. This unit (Thl in Fig. 24) capped a series of nonmarine sediments that filled interior basins in this region at the close of Miocene tectonic extension (e.g., Tm). The highest outcrops of this limestone represent the level of the interior basin fill just before arrival of and incision by the Colorado River (Lucchitta, 1972; Longwell, 1936; Blair and Armstrong, 1979). The highest limestone outcrops can be traced by eye as light-colored mesas 6–10 km to the northeast and above our position, at an elevation of ~700 m (2300 ft) asl. They outline a bathtub ring along the southeast side of the valley that laps against dark Miocene volcanic rocks. The high limestone caps a sequence of mudstones, gypsum, more limestone, and unexposed halite deposits. This upper Miocene fill sequence evidently occupied a bolson not unlike many others in the modern Basin and Range province. Spencer et al. (2001) dated a tuff in the Hualapai Limestone just beyond the field of view to the southeast, a few tens of meters below the highest outcrops. The date of 5.97 ± 0.07 Ma constrains the timing of the subsequent arrival of the river into the Basin and Range province. The highest exposures of the Hualapai occupy a similar elevation in the next valley to the east and over the top of the next adjacent range to the east. At the base of the Grand Wash Cliffs the limestone top has been uplifted along a fault to as high as 880 m (2900 ft) asl. The Hualapai Limestone was suspected to be marine by Blair (1978), and nonmarine but deposited near sea level by Lucchitta (1979) and Lucchitta et al. (2001). Strontium isotope measurements show that the Hualapai is nonmarine and isotopically unlike the modern Colorado River, but the Sr values are consistent with a preriver origin from sources affected by the Precambrian bedrock (Spencer and Patchett, 1997; Patchett and Spencer, 2001). Oxygen isotopic values indicate that the Hualapai Limestone was mostly nonevaporative, and its basin of deposition therefore saw considerable through-flow of water (Faulds et al., 2001). The oldest direct evidence of the river arriving in the Basin and Range province in the Lake Mead region are rare roundstone pebbles intercalated in a sequence a few meters thick of sand, silt, and clay that conformably overlies the Hualapai Limestone near the mouth of the Grand Canyon. Subsequent Colorado River sands and gravels are incised into the Hualapai Limestone and below it at a variety of levels from a few meters to hundreds of meters (Howard and Bohannon, 2001).

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Figure 24. Air-photo map and conceptional cross-sectional sketch of the northeast side of Detrital Valley. Detrital Wash on the left drains northward to Lake Mead. The geologic units, in upward stratigraphic succession, are as follows: Tv—middle Miocene mafic volcanic rocks and ash-flow tuff; Tm—upper Miocene pink mudstone, limestone, and gypsum; Thl—upper part of the Hualapai Limestone (upper Miocene), which outcrops as a bathtublike ring buttressed against the volcanic highland; Tr—riverlaid sand and gravel of the ancestral Colorado River, inset below the Hualapai Limestone (Pliocene?); and Taf—mesa-forming fanglomerate graded onto the Tr deposits and capped by a thick, resistant calcic soil (Pliocene?). Younger alluvial-fan deposits are not labeled.

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One such packet of fluvial gravel and sand can be seen in front of and inset into the Hualapai Limestone exposures in the field of view, on the southwest flank of Detrital Valley. The undated river deposit (Tr) here is horizontal and lies at elevations between 520 and 620 m, 80 m below the floor of Hualapai Limestone. Its 100m thickness represents an aggradational phase in the early river history. Whether the deposit reflects the same aggradation as the alluvium of Bullhead City is not yet resolved (Beard et al., 2005). Northward toward Lake Mead are other deposits at successively lower elevation. Because exposures concentrate at two or more distinct levels (Mel Kuntz, 1997, personal commun.) we suspect they are inset fills, younger than the 100-m-thick section before you; however it is also possible that they represent strath terraces cut into lower and deeper parts of an early, thick Bullhead-type aggradation. Lee (1908) proposed that the river initially followed Detrital Valley far southward to Sacramento Valley and the Needles area, but there is no evidence to support this or any other course greatly different from the modern one. The existence of river gravels at high levels in this and other valleys and considerably away from the modern west-flowing channel, however, suggests that at times

of aggradation the river wandered extensively on braidplains away from its most direct westward route, which was directly across these N-striking valleys. Graded over the top of the river gravels are remnants of locally derived alluvial-fans capped by mesa-forming resistant, meters-thick soil carbonate horizons (Taf). The thick soil is consistent with a Pliocene age. As with the old fans discussed earlier in the trip near Golden Shores, these fans may record progradation over abandoned river deposits subsequent to incision of the river and erosion of river deposits. How many incision-aggradation sequences the river experienced in the Lake Mead area is yet to be determined. One important clue may come, however, from dating in progress by Ari Matmon of abandoned river potholes near Hoover Dam, at Stop 3.3. Directions and Highlights en Route to Stop 3.3 Return southward up the Temple Bar Road and reset the odometer as you turn north onto U.S. 93 toward Hoover Dam (Fig. 22). In two mi you will leave the high detritus of Detrital Valley through Householder Pass and see the deep canyon of the Colorado River. This pass is barely higher than the Hualapai

Birth of the lower Colorado River Limestone and the inferred initial path of the river, yet there is no evidence that the early river took this shortcut southwestward. Instead the river coursed westward and incised precipitous Boulder Canyon through a high ridge of the Black Mountains 35 km north of here, then flowed westward through Boulder Basin before exiting southward into Black Canyon, the site of Hoover Dam, and continuing into the canyon in the foreground. This surprising incision, together with observations that old river gravels near Boulder Canyon are folded (Longwell, 1936; Anderson, 2003), suggest that the northern Black Mountains—and perhaps Boulder Canyon—have been uplifted since the river first established its course (Howard and Bohannon, 2001). The views into Black Canyon give the viewer an appreciation of the magnitude of the river’s incision. The drive northward along the east side of the canyon traverses basin-fill deposits— fanglomerate and megabreccia (avalanche) deposits—that have been deeply incised by the river (Anderson, 1978). In 8–9 mi the road cuts through a ca. 5-Ma basalt flow (Anderson et al., 1972) that slopes westward on a gradient consistent with deposition into a shallow pre–Colorado River or early Colorado River valley (Howard and Bohannon, 2001). Beyond the Temple Bar Road 22–23 mi you will begin to see evidence of a massive construction project for a highway to bypass Hoover Dam and reestablish Hwy 93 as an important commercial route. The bypass will culminate in a 900-ft-high bridge across Black Canyon. In another mile, just before beginning to descend switchbacks to Hoover Dam, turn right into the parking area for an overview of Hoover Dam, and park. Note: Visiting Stop 3.3 requires pre-approval from the Hoover Dam Police. Stop 3.3: Potholes Just too Pretty to Repair A small and unique exposure of river-sculpted potholes and overlying fluvial gravel deposited by the early Colorado River lie stranded on the rim of Black Canyon 275 m above the river bed (Fig. 25). This exposure records a time when the river sculpted and exposed this site and partly buried it under gravels before ultimately abandoning this course and further incising the deep canyon in which Hoover Dam was built. Ransome (1923) recognized the significance of these highlevel potholes as evidence of the river’s course before the adjacent river canyon had been fully cut. In what may be one of the earliest practical applications of paleoseismology, he observed that the potholes cut across faults in the volcanic bedrock, and inferred from the depth of subsequent canyon incision that the potholes are old and that such faults near the dam site therefore must be ancient enough to pose no earthquake risk to a future dam. The dam was built, and time has confirmed his inference so far. The river-sculpted pothole surfaces record an incision depth on the order of 200 m below the earliest Pliocene preriver topographic surface, which is recorded by basin fill including the 5 ± 0.4 Ma Fortification basalt flow a few miles to the northwest of our position (Anderson et al., 1972; Damon et al., 1978; Mills, 1994) and by a paleogroundwater table mapped by Anderson (1969). Since pothole formation, the river has incised another

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Figure 25. River-sculpted potholes in dacite, overlooking Hoover Dam.

275 m. As much as 25 m of vertical relief can be seen on riversculpted bedrock and in river gravels that line and overlie the bedrock paleochannel (Howard et al., 2004). Cemented gravel at the same level is exposed in a cut for the new bypass highway ~1 km downstream. The sculpted rock gorge filled by river sediment here may have a modern analog under Hoover Dam. During excavation for the dam, the discovery of sawn plank buried in river fill 15 m below the low river level at a bedrock bench led Berkey (1935) to conclude that the river bed was recently reworked to that depth. A narrow inner gorge notched another 22 m deeper was lined with potholes and vertically fluted bedrock surfaces, features that led Berkey (1935) to conclude that the inner gorge was cut by pothole incision. REFERENCES CITED Agenbroad, L.D., Mead, J.I., and Reynolds, R.E., 1992, Mammoths in the Colorado River corridor, in Reynolds, R.E., compiler, Old routes to the Colorado: San Bernardino County Museum Special Publication 92-2, p. 104–106. Anderson, R.E., 1969, Notes on the geology and paleohydrology of the Boulder City Pluton, southern Nevada: U.S. Geological Survey Professional Paper 650-B, p. 35–40. Anderson, R.E., 1971, Thin-skinned distension in Tertiary rocks of southeastern Nevada: Geological Society of America Bulletin, v. 82, p. 43–58.

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Anderson, R.E., 1978, Geologic map of the Black Canyon 15-minute quadrangle, Mohave County, Arizona, and Clark County, Nevada: U.S. Geological Survey Map GQ-1394, scale 1:62,000. Anderson, R.E., 2003, Geologic map of the Callville quadrangle, Nevada: Nevada Bureau of Mines and Geology Map 139, scale 1:24,000. Anderson, R.E., Longwell, C.R., Armstrong, R.L., and Marvin, R.F., 1972, Significance of K-Ar ages of Tertiary rocks from the Lake Mead region, Nevada-Arizona: Geological Society of America Bulletin, v. 83, p. 273–288. Bahmoier, H.F., 1950, Construction engineering problems at Davis Dam: Advance copy of paper presented at the 1950 Spring Meeting of the American Society of Civil Engineers, Los Angeles, California, April 26–29. Beard, L.S., Felger, T.J., House, P.K., and Howard, K.A., 2005, Transition from basins to through-flowing drainage of the Colorado River, lower Lake Mead region, Nevada and Arizona, in Reheis, M.C., Geologic and biotic perspectives on Late Cenozoic drainage history of the southwestern Great Basin and lower Colorado River region: Conference Abstracts, Geologic and Biotic Perspectives, Zzyzx, California, April 12–15: U.S. Geological Survey Open-File Report (in press). Bell, J.W., Ku, T.-L., and Kukla, G.J., 1978, The Chemehuevi Formation of Nevada, Arizona, and California: An examination of its distribution, facies, and age: Geological Society of America Abstracts with Programs, v. 10, no. 3, p. 95. Berkey, C.P., 1935, Geology of Boulder and Norris Dam sites: Civil Engineering, v. 5, p. 24–28. Blackwelder, E., 1933, Terraces along the lower course of the Colorado River [abs.]: Proceedings of the Geological Society of America, p. 66. Blackwelder, E., 1934, Origin of the Colorado River: Geological Society of America Bulletin, v. 231, p. 551–566. Blair, J.L., 1996, Drastic modification of the depositional style of the lower Colorado River in late Pleistocene time: Evidence from fine-grained strata in the Lake Mohave area, Nevada/Arizona [M.S. thesis]: Nashville, Tennessee, Vanderbilt University, 138 p. Blair, W.N., 1978, Gulf of California in Lake Mead area of Arizona and Nevada during late Miocene time: AAPG Bulletin, v. 62, p. 1159–1170. Blair, W.N., and Armstrong, A.K., 1979, Hualapai Limestone member of the Muddy Creek Formation: The youngest deposit predating the Grand Canyon, southeastern Nevada and northwestern Arizona: U.S. Geological Survey Professional Paper 1111, 14 p. Buising, A.V., 1990, The Bouse Formation and bracketing units, southeastern California and western Arizona: Implications for the evolution of the proto-Gulf of California and the lower Colorado River: Journal of Geophysical Research, v. 95, p. 20,111–20,132. Campbell, E.A., and John, B.E., 1996, Constraints on extension-related plutonism from modeling of the Colorado River gravity high: Geological Society of America Bulletin, v. 108, p. 1242–1255, doi: 10.1130/00167606(1996)108<1242:COERPF>2.3.CO;2. Carr, W.J., 1991, A contribution to the structural history of the Vidal-Parker region, California and Arizona: U.S. Geological Survey Professional Paper 1430, 40 p. Damon, P.E., Shafiqullah, M., and Scarborough, R.B., 1978, Revised chronology for critical stages in the evolution of the lower Colorado River: Geological Society of America Abstracts with Programs, v. 10, no. 3, p. 101. Dickey, D.D., Carr, W.J., and Bull, W.B., 1980, Geologic map of the Parker NW, Parker, and parts of the Whipple Mountains SW and Whipple Wash quadrangles: U.S. Geological Survey Miscellaneous Investigations Series Map I-1124, scale 1:24,000. Faulds, J.E., 1996a, Geologic map of the Mt. Davis Quadrangle, Nevada and Arizona: Nevada Bureau of Mines and Geology Map 105, 1:24,000 scale. Faulds, J.E., 1996b, Geologic map of the Fire Mountain Quadrangle, Nevada and Arizona, Nevada Bureau of Mines and Geology Map 106, 1:24,000 scale. Faulds, J.E., and House, P.K., 2000, Geology of the Laughlin area, Clark County, Nevada, in Faulds, J.E., House, P.K., Shevenell, L., and Ramelli, A.R., 2000, Geology and natural hazard assessment of the Laughlin area, Clark County, Nevada: Nevada Bureau of Mines and Geology Open-File Report 2000-6, p. 1.1–1.56. Faulds, J.E., Wallace, M.A., Gonzales, L.A., and Heizler, M.T., 2001, Depositional environment and paleogeographic implications of the late Miocene Hualapai Limestone, northwestern Arizona and southern Nevada, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 81–87.

Faulds, J.E., Gonzalez, L.A., Perkins, M.E., House, P.K., Pearthree, P.A., Castor, S.B., and Patchett, J.P., 2002, Late Miocene–early Pliocene transition from lacustrine to fluvial deposition: Inception of the Lower Colorado River in southern Nevada and northwest Arizona: Geological Society of America Abstracts with Programs, v. 34, no. 4, p. A-60. Faulds, J.E., House, P.K., Pearthree, P.A., Bell, J.W., and Ramelli, A.R., 2004, Preliminary geologic map of the Davis Dam quadrangle and eastern part of the Bridge Canyon quadrangle, Clark County, Nevada and Mohave County, Arizona: Nevada Bureau of Mines and Geology Open-File report 03-5, 1 sheet, scale 1:24,000. Glancy, P.A., and Harmsen, L., 1975, A hydrologic assessment of the September 14, 1974 flood in Eldorado Canyon, Nevada: U.S. Geological Survey Professional Paper 930, 69 p. Gray, F., Jachens, R.C., Miller, R.J., Turner, R.L., Knepper, D.H., Pitkin, J.A., Keith, W.J., Mariano, J., and Jones, S.L., 1990, Mineral resources of the Warm Springs Wilderness Study Area, Mohave County, Arizona: U.S. Geological Survey Bulletin 1737, 20 p. Gross, E.L., Patchett, P.J., Dallegge, T.A., and Spencer, J.E., 2001, The Colorado River system and Neogene sedimentary formations along its course: Apparent Sr isotopic connections: Journal of Geology, v. 109, p. 449–461, doi: 10.1086/320793. House, P.K., Pearthree, P.A., Bell, J.W., Ramelli, A.R., and Faulds, J.E., 2002, New stratigraphic evidence for the Late Cenozoic inception and subsequent alluvial history of the lower Colorado River from near Laughlin, Nevada: Geological Society of America Abstracts with Programs, v. 34, no. 4, p. A-60. House, P.K., Howard, K.A., Bell, J.W., and Pearthree, P.A., 2004a, Preliminary geologic map of the Arizona and Nevada parts of the Mt. Manchester Quadrangle: Nevada Bureau of Mines and Geology Open-file Report 0404, scale 1:24,000. House, P.K., Pearthree, P.A., Faulds, J.E., and Bell, J.W., 2004b, Alluvial and lacustrine stratigraphic evidence for the late Neogene inception and early evolution of the Lower Colorado River in the vicinity of Pyramid Canyon, Nevada–Arizona: Geological Society of America Abstracts with Programs, v. 36, no. 5, p. 550. Howard, K.A., and Bohannon, R.G., 2001, Lower Colorado River; Framework, Neogene deposits, incision, and evolution, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 101–105. Howard, K.A., and John, B.E., 1987, Crustal extension along a rooted system of imbricate low-angle faults, Colorado River extensional corridor, California and Arizona, in Coward, M.P., Dewey, J.F., and Hancock, P.L., eds., Continental extensional tectonics: London, Geological Society Special Publication 28, p. 299–311. Howard, K.A., and John, B.E., 1997, Fault-related folding during extension: Plunging basement-cored folds in the Basin and Range: Geology, v. 25, p. 223–226, doi: 10.1130/0091-7613(1997)025<0223:FRFDEP>2.3.CO;2. Howard, K.A., Faulds, J.E., Beard, L.S., and Kunk, M.J., 2000, Reverse-drag folding across the path of the antecedent early Pliocene Colorado River below the mouth of the Grand Canyon: Geological Society of America Abstracts with Programs, v. 32, no. 7, p. 41. Howard, K.A., Lundstrom, S.C., and Matmon, A., 2004, Ancestral Colorado River potholes high above Hoover Dam: Geological Society of America Abstracts with Programs, v. 36, no. 5, p. 515. Ives, J.C., 1861, Report upon the Colorado River of the West: Washington, 36th Congress 1st session, Senate, Government Printing Office, in McKinney, K.C., 2002, ed., Digital archive-report upon the Colorado River of the West explored in 1857 and 1858 by Lieutenant Joseph C. Ives, geological report with maps by John S. Newberry: U.S. Geological Survey Open-file Report 02-25, 154 p., CD-ROM. John, B.E., 1987, Geometry and evolution of a mid-crustal extensional fault system: Chemehuevi Mountains, southeastern California, in Coward, M.P., Dewey, J.F., and Hancock, P.L., eds., Continental extensional tectonics: London, Geological Society Special Paper. 28, p. 313–335. Johnson, N.M., Officer, C.B., Opdyke, N.D., Woodward, G.D., Zeitler, P.K., and Lindsay, E.H., 1983, Rates of late Cenozoic tectonism in the Vallecito-Fish Creek basin, western Imperial Valley, California: Geology, v. 11, p. 664–667, doi: 10.1130/0091-7613(1983)11<664:ROLCTI>2.0.CO;2. Ku, T.L., Bull, W.B., Freeman, S.T., and Knauss, K.G., 1979, Th230–U234 dating of pedogenic carbonates in gravelly desert soils of Vidal Valley, southeastern California: Geological Society of America Bulletin, v. 90, p. 1063– 1073, doi: 10.1130/0016-7606(1979)90<1063:TDOPCI>2.0.CO;2.

Birth of the lower Colorado River Kukla, G.J., 1975, Preliminary report on magnetostratigraphic study of sediments near Blythe and Parker Valley, California and Arizona: Appendix 2.5B, Early Site Review Report (archived), Sundesert Nuclear Power Project, 29 p. Lee, G.K., and Bell, J.W., 1975, Depositional and geomorphic history of the lower Colorado River: Appendix 2.5D, Early Site Review Report (archived), Sundesert Nuclear Power Project, 25 p. Lee, W.T., 1908, Geologic reconnaissance of a part of western Arizona: U.S. Geological Survey Bulletin 252, p. 41–45. Longwell, C.R., 1936, Geology of the Boulder Reservoir floor, ArizonaNevada: Geological Society of America Bulletin, v. 47, p. 1393–1476. Longwell, C.R., 1947, How old is the Colorado River?: American Journal of Science, v. 244, p. 817–835. Longwell, C.R., 1963, Reconnaissance geology between Lake Mead and Davis Dam, Arizona-Nevada: U.S. Geological Survey Professional Paper 374-E, 51 p. Lucchitta, I., 1972, Early history of the Colorado River in the Basin and Range province: Geological Society of America Bulletin, v. 83, p. 1933–1948. Lucchitta, I., 1979, Late Cenozoic uplift of the southwestern Colorado River region: Tectonophysics, v. 61, p. 63–95, doi: 10.1016/0040-1951(79)90292-0. Lucchitta, I., 1998, The upper Miocene Bouse Formation as an indicator for late Cenozoic uplift of the Colorado Plateau: Geological Society of America Abstracts with Programs, v. 30, p. A-14. Lucchitta, I., McDougall, K., Metzger, D.G., Morgan, P., Smith, G.R., and Chernoff, B., 2001, The Bouse Formation and post-Miocene uplift of the Colorado Plateau, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 173–178. Lundstrom, S.C., Mahan, S.A., Hudson, M.R., and Paces, J.B., 2000, Evidence for extreme Pleistocene floods of the lower Colorado River, in AMQUA 2000: Fayetteville, Arkansas, Program and Abstract of the 16th Biennial Meeting, May 22–24, p. 81. Lundstrom, S.C., Mahan, S.A., Paces, J.B., and Hudson, M.R., 2004, Late Pleistocene aggradation and incision of the lower Colorado River downstream of the Grand Canyon: Geological Society of America Abstracts with Programs, v. 36, no. 5, p. 550. Meek, N., and Douglass, J., 2001, Lake overflow: An alternative hypothesis for Grand Canyon incision and development of the Colorado River, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 199–206. Merriam, R., and Bischoff, J.L., 1975, Bishop ash: A widespread volcanic ash extended to southern California: Journal of Sedimentary Petrology, v. 45, p. 207–211. Metzger, D.G., 1968, The Bouse Formation (Pliocene) of the Parker-BlytheCibola area, Arizona and California, in Geological Survey Research 1968: U.S. Geological Survey Professional Paper 600-D, p. D126–D136. Metzger, D.G., and Loeltz, O.J., 1973, Geohydrology of the Needles area, Arizona, California, and Nevada: U.S. Geological Survey Professional Paper 486-J, 54 p. Metzger, D.G., Loeltz, O.J., and Irelan, B., 1973, Geohydrology of the ParkerBlythe-Cibola area, Arizona and California: U.S. Geological Survey Professional Paper 486-G, 130 p. Miller, M.J., and John, B.E., 1999, Sedimentation patterns support seismogenic low-angle normal faulting, southeastern California and western Arizona: Geological Society of America Bulletin, v. 111, p. 1350–1370, doi: 10.1130/0016-7606(1999)111<1350:SPSSLA>2.3.CO;2. Mills, J.G., 1994, Geologic map of the Hoover Dam quadrangle, Arizona and Nevada: Nevada Bureau of Mines and Geology, Map 102, scale 1:24,000. Morgan, L.A., and McIntosh, W.C., 2005, Timing and development of the Heise volcanic field, Snake River Plain, Idaho, western USA: Geological Society of America Bulletin, v. 117, p. 288–306, doi: 10.1130/B25519.1. Olmstead, F.H., Loeltz, O.J., and Irelan, B., 1973, Geohydrology of the Yuma area, Arizona and California: U.S. Geological Survey Professional Paper 486-H, 154 p. Patchett, P.J., and Spencer, J.E., 2001, Application of Sr isotopes to the hydrology of the Colorado River system waters and potentially related Neogene sedimentary formations, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 167–171.

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Poulson, S.R., and John, B.E., 2003, Stable isotope and trace element geochemistry of the basal Bouse Formation carbonate, southwestern United States: Implications for the Pliocene uplift history of the Colorado Plateau: Geological Society of America Bulletin, v. 115, p. 434–444, doi: 10.1130/0016-7606(2003)115<0434:SIATEG>2.0.CO;2. Pearthree, P.A., and House, P.K., 2004, Digital geologic map of the Davis Dam southeast quadrangle, Mohave County, Arizona, and Clark County, Nevada: Arizona Geological Survey Digital Geologic Map DGM-45, scale 1:24,000. Pearthree, P.A., Menges, C.M., and Mayer, L., 1983, Distribution, recurrence, and possible tectonic implications of late Quaternary faulting in Arizona: Tucson, Arizona Bureau of Geology and Mineral Technology Open-File Report 83-20, 51 p. Purcell, C., and Miller, D.G., 1980, Grabens along the lower Colorado River, California and Arizona, in Fife, D.L, and Brown, A.R., eds., Geology and mineral wealth of the California desert: Santa Ana, California, South Coast Geological Society, 555 p. Ransome, F.L., 1923, Ancient high-level potholes near the Colorado River: Science, New Series, v. 57, p. 593. Sarna-Wojcicki, A.M., Lajoie, K.R., Meyer, C.E., Adam, D.P., and Rieck, H.J., 1991, Tephrochronologic correlation of upper Neogene sediments along the Pacific margin, conterminous United States, in Morrison, R.B., ed., Quaternary nonglacial geology; Conterminous U.S.: Boulder, Colorado, Geological Society of America, The Geology of North America, v. K-2, p. 117–140. Scarborough, R., 2001, Neogene development of the Little Colorado River Valley and Eastern Grand Canyon: Field evidence for an overtopping hypothesis, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 215–222. Simpson, R.W., Howard, K.A., Jachens, R.C., and Mariano, J., 1990, A positive gravity anomaly along the Colorado River extensional corridor: Evidence for new crustal material: Eos (Transactions, American Geophysical Union), v. 71, p. 1594. Spencer, J.E., 1985, Miocene low-angle normal faulting and dike emplacement, Homer Mountain and surrounding areas, southeastern California and southernmost Nevada: Geological Society of America Bulletin, v. 96, p. 1140– 1155, doi: 10.1130/0016-7606(1985)96<1140:MLNFAD>2.0.CO;2. Spencer, J.E., and Patchett, P.J., 1997, Sr isotope evidence for a lacustrine origin for the upper Miocene to Pliocene Bouse Formation, lower Colorado River trough, and implications for timing of Colorado Plateau uplift: Geological Society of America Bulletin, v. 109, p. 767–778, doi: 10.1130/ 0016-7606(1997)109<0767:SIEFAL>2.3.CO;2. Spencer, J.E., and Pearthree, P.A., 2001, Headward erosion versus closed-basin spillover as alternative causes of Neogene capture of the ancestral Colorado River by the Gulf of California, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 215–222. Spencer, J.E., and Reynolds, S.J., 1989, Middle Tertiary tectonics of Arizona and adjacent areas, in Jenney, J.P., and Reynolds, S.J., eds., Geologic evolution of Arizona: Arizona Geological Society Digest 17, p. 539–574. Spencer, J.E., Peters, L., McIntosh, W.C., and Patchett, P.J., 2001, 40A/39Ar geochronology of the Hualapai Limestone and Bouse Formation and implications for the age of the lower Colorado River, in Young, R.A., and Spamer, E.E., eds., The Colorado River: Origin and evolution: Grand Canyon, Arizona, Grand Canyon Association Monograph 12, p. 89–91. Spencer, J.E., Pearthree, P.A., Patchett, J., and House, P.K., 2005, Evidence for a lacustrine origin for the lower Pliocene Bouse Formation, lower Colorado River Valley, in Reheis, M.C., Geologic and biotic perspectives on Late Cenozoic drainage history of the southwestern Great Basin and lower Colorado River region: Conference Abstracts, Geologic and Biotic Perspectives, Zzyzx, California, April 12–15: U.S. Geological Survey Open-File Report (in press). U.S. Bureau of Reclamation, 1955, Technical record of design and construction, Davis Dam, Chapter II: Geology: Washington, D.C., U.S. Bureau of Reclamation, p. 9–27. Whitney, J.W., 1996, Evidence of Quaternary faulting in Las Vegas Wash, Clark County, Nevada, in dePolo, C.M., ed., Proceedings of a Conference on Seismic Hazards in the Las Vegas Region: Nevada Bureau of Mines and Geology, Open-File Report 98-6, p. 76.

Printed in the USA

Geological Society of America Field Guide 6 2005

Development of Miocene faults and basins in the Lake Mead region: A tribute to Ernie Anderson and a review of new research on basins Melissa Lamb Geology Department, OWS 153, University of St. Thomas, 2115 Summit Ave, St. Paul, Minnesota 55105, USA Paul J. Umhoefer Department of Geology, 4099, Northern Arizona University, Flagstaff, Arizona 86011, USA Ernie Anderson U.S. Geological Survey, Box 347, Kernville, California 93238, USA L. Sue Beard U.S. Geological Survey, 2255 N. Gemini Drive, Flagstaff, Arizona 86001, USA Thomas Hickson Geology Department, OWS 153, University of St. Thomas, 2115 Summit Ave, St. Paul, Minnesota 55105, USA K. Luke Martin Department of Geology, 4099, Northern Arizona University, Flagstaff, Arizona 86011, USA

ABSTRACT The purpose of this field trip is to provide an overview of Miocene basin development in the Lake Mead region, demonstrate how basin-fill deposits reflect tectonic activity on a variety of structures, and highlight the work of Ernie Anderson in this region. The Basin and Range province is superb for the study of major normal and strike-slip fault systems that accommodate large-magnitude extension. Within this province, the Lake Mead region provides exceptional exposures of synextensional Miocene basins and faults and is a transition zone between predominantly half-graben–style basins and ranges to the north and the highly extended Colorado River Extensional Corridor to the south. The region also embraces a change from thick Phanerozoic sedimentary rocks in the north to Precambrian crystalline basement rocks overlain by late Tertiary volcanic rocks in the south. The early Paleozoic “Cordilleran hingeline” and the southeast margin of east-directed Mesozoic thrusting are also within this transition zone, but the area contains a strong overprint of late Tertiary tectonism. This overprint is strongest near the intersection between two major strike-slip fault systems: the right-lateral Las Vegas Valley shear zone and the left-lateral Lake Mead fault system. Miocene sedimentary rocks record the onset of major

E-mails: [email protected]; [email protected]; [email protected]; [email protected]; [email protected]; [email protected]. Lamb, M., Umhoefer, P.J., Anderson, E., Beard, L.S., Hickson, T., and Martin, K.L., 2005, Development of Miocene faults and basins in the Lake Mead region: A tribute to Ernie Anderson and a review of new research on basins, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 389–418, doi: 10.1130/2005.fld006(18). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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extension and the development of numerous, complex structures. Details of the extension and the resulting complexity are not fully understood and many issues remain unresolved. The relative importance of normal versus strike-slip faults is debated as are the details of how and why faults develop through time. Keywords: Miocene extension, Basin and Range, normal fault, sedimentary basin, strike-slip fault.

INTRODUCTION AND OBJECTIVES Continental extension is a fundamental geologic process that results in crustal thinning and basin development. Large amounts of extension or rifting of the continents can ultimately lead to the development of passive continental margins and ocean basins. Many of the world’s largest petroleum deposits accumulate in rift basins or at passive margins. These observations highlight the need to understand fully the geometry, kinematics, and dynamics of extensional processes and the major features of extensional fault systems and related basins. Most young and active rift–passive margin systems are difficult to study because a large portion of them has subsided below sea level. The well-exposed Basin and Range province of the western U.S. is a superb natural laboratory for the study of the development of major normal and strike-slip fault systems that accommodate large-magnitude crustal extension. The Lake Mead region, Nevada and Arizona (Figs. 1–4), is part of the Basin and Range province and provides exceptional exposures of synextensional Miocene basins and faults. Within these basins, there are world-class exposures of clastic, carbonate, and evaporite strata that annually attract tens of thousands of tourists and academic and industry scientists. The purpose of this field trip is to provide participants an overview of Miocene basin development in the Lake Mead region and show how the basin-fill deposits reflect a response to tectonic activity on a wide variety of extensional, strike-slip, and contractional structures. The faults and basins lie within a structurally complex zone formed by the interaction of the right-lateral Las Vegas Valley shear zone and the left-lateral Lake Mead fault system (Fig. 3). Within this zone, depocenters shifted positions, changed shape, were rotated and/or inverted and are thought to have evolved from broad to restricted in response to strike-slip and normal fault displacements that ranged to tens of kilometers. Syntectonic sedimentation in this area shaped an extraordinarily complex geology that includes ongoing neotectonics with hazards significance. Over the three days of this trip, we will progress upward through the Miocene stratigraphic section along a west to east transect north of Lake Mead, from the Frenchman Mountain area near Las Vegas to the Overton Arm of Lake Mead (Figs. 3–5). We will examine the evolution of Miocene faults and related basins as we present the history of development of ideas, results of ongoing studies, and needs for future studies to resolve existing problems at each stratigraphic level. In particular, we will focus on two types of studies: (1) critical framework studies that have resulted in much debate about the relation of extensional, strike-

slip, and contractional structures within this belt; and (2) ongoing studies that are providing critical new data, such as structural setting of individual basins and age and magnitude of extension and contraction, that can help constrain tectonic reconstructions and critique or expand existing models. We will evaluate the challenge presented to the well-established published stratigraphy by new field and radiometric age data. Also, the existing paradigm that Miocene depocenters evolved from broad basins to smaller subbasins in response to increasing displacements on the major fault systems will be evaluated in the light of recent studies. The final important objective of this field trip is to highlight the work of Ernie Anderson in contributing to understanding the Lake Mead region in general and specifically to the extremely complex area where the Las Vegas Valley shear zone and the Lake Mead fault system interact. We will visit critical outcrops that highlight relationships deemed important by Anderson and discuss how future work might make progress toward resolving some of the controversial issues that have been highlighted in published reports. GEOLOGIC SETTING We define the Lake Mead region as extending from the Las Vegas Valley on the west to the Colorado Plateau on the east and from the Virgin River depression (I-15 corridor) on the north to approximately the south shore of Lake Mead on the south (Figs. 1 and 3). This region embraces a major gravity step reflecting a transition from high elevations (ca. 1.5 km) to the north to low elevations (ca. 0.5 km) to the south (Eaton et al., 1978). This transition zone is situated between an area of predominantly halfgraben–style basins and intervening ranges—classic Basin and Range structure—to the north and the highly extended Colorado River Extensional Corridor to the south. The Lake Mead region is the eastern part of a classic large-magnitude extensional domain reaching from the Colorado Plateau on the east to Death Valley and the eastern edge of the Sierra Nevada on the west (e.g., Wernicke et al., 1988; the Central Basin and Range area within Fig. 1A). This east-west–trending extensional corridor is thought to have undergone 250–300 km of extension in the late Cenozoic from ca. 16 Ma to the present (Wernicke et al., 1988; Snow and Wernicke, 2000). This field trip focuses on a small part of the Lake Mead region reaching from Frenchman Mountain on the west to the Overton Arm of Lake Mead on the east. We refer to this limited area as the Frenchman-Overton corridor (Fig. 1B). The Frenchman-Overton

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Figure 1. (A) Map of the southwestern United States showing physiographic provinces. Inset box shows location of 1B. Modified from Wernicke et al. (1988) and Stewart (1998). (B) Map showing physiographic features, major extensional structures, and structural domains of the eastern portion of the Central Basin and Range. Shaded outlines indicate ranges; white areas indicate basins. Heavy dashed line indicates FrenchmanOverton Corridor and the focus of the field trip. Modified from Axen et al. (1993), Price (1997), and Spencer and Reynolds (1989). Timing constraints from Axen et al. (1990, 1993) (Morman Mountain domain), Duebendorfer et al. (1998) (Lake Mead domain), and Spencer and Reynolds (1989) (Whipple domain). BMAZ—Black Mountain accommodation zone; CCD—Castle Cliff detachment; CMF—Cerbat Mountains fault; DMF—Dupont Mountain fault; GWF—Grand Wash fault; HSF—Horse Spring fault; KSWF—Kane Springs Wash fault; LBR—Lost Basin Range; LMFS—Lake Mead fault system; LVVSZ—Las Vegas Valley shear zone; MSF—Mountain Spring fault; PFZ—Pahranagat fault zone; SVWHD—South Virgin-White Hills detachment; TSD—Tule Springs detachment; WBRBD—Whipple-Buckskin-Rawhide-Bullard detachment; WRF—Wheeler Ridge fault.

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Figure 3. Generalized geologic map of the Lake Mead region (modified from Campagna and Aydin, 1994; Duebendorfer and Sharp, 1998; Duebendorfer et al., 1998). Stars indicate general area visited each day. BRF—Bitter Ridge fault; BSVF—Bitter Spring Valley fault; CB—Callville Bay; CL—Cleopatra Lobe of the HamblinCleopatra Volcano; FM—Frenchman Mountain; GBF—Gold Butte fault; GH—Gale Hills; HBF—Hamblin Bay fault; HL—Hamblin Lobe of the Hamblin-Cleopatra Volcano; HSF—Hen Spring fault; LR—East and West Longwell Ridges; LRF—Lime Ridge fault; NBM—Northern Black Mountains; RSF—Rogers Spring fault; SVM— south Virgin Mountains; WH—White Hills. The Lake Mead fault system (LMFS) includes the BSVF, HBF, RSF, HSF and BRF.

Day 3 Lake Mead

Overton Arm area

Day 2

Virgin Mtns area

Grand Wash Trough

Paleozoic sedimentary rocks Proterozoic crystalline rocks

Younger Tertiary deposits Horse Spring Formation

Tertiary volcanic rocks

Tertiary intrusive rocks

Mesozoic sedimentary rocks Surficial Deposits

Figure 4. Generalized geologic map of the Lake Mead 1:100,000 quadrangle, modified from Beard et al. (2006). Boxes indicate locations of field trip stops for each day.

Areas noted by yellow boxes indicate locations for Cenozoic sections in Fig. 5

GENERALIZED GEOLOGIC MAP OF THE LAKE MEAD 30 X 60 QUADRANGLE, NEVADA - ARIZONA

Day 1

Frenchman Mtn

Frenchman MtnMuddy Mtns area

Muddy Mountains

Development of Miocene faults and basins in the Lake Mead region 393

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Hoover Dam volcanic and plutonic rocks (13.9-12.7 Ma)

Hamblin-Cleopatra volcano (10.1-13.1 Ma)

Callville Mesa volcanics (8.5-10.6 Ma)

Fortification Hill basalt (5.8 Ma)

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Thumb Member (14.2-16.2 Ma)

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Thumb Member conglomerate (13.93 Ma)

Tob

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Muddy Creek Fm (8 Ma)

'Hualapai Limestone’ (6.0 Ma)

Twin Springs sandstone/gypsum deposits (13.2 Ma)

Tmc

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Post-extension

CENOZOIC STRATIGRAPHIC CORRELATION OF LAKE MEAD GEOLOGY

Figure 5. Cenozoic stratigraphy from the Lake Mead region. See Figure 4 for column locations.

Rainbow Gardens member

River Mtn Tvr volcanics Wilson Ridge (12.5-13.2 Ma) Twr pluton (12.6 Bitter Ridge Thd -13.5 Ma) Thb Limestone Mt. Davis Member volcanics Tmd (13.0-13.5 Ma) (13.1-15.0 Ma) Boulder Wash Tbw volcanics (14.2 Ma) Thumb Member Tpm Tht (13.5-17.2 Ma) Patsy Mine volcanics (15.6-14.2 Ma)

Red Sandstone Unit (11.9 - 10.6 Ma)

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TERTIARY

QUATERNARY

Holocene

Surficial deposits

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Snap Point basalt (8.8 Ma)

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Rainbow Gardens Member

Thumb Member (15.3 Ma)

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Tsp

Limestone of Grand Wash Trough Tlgw (7.4 - 10.9 Ma)

Tbgw

Grand Wash Trough

394 M. Lamb et al.

Horse Spring Fm

Development of Miocene faults and basins in the Lake Mead region corridor approximately traces the early Paleozoic “Cordilleran hingeline” (or Wasatch line) and the southeast margin of the area affected by east-directed Mesozoic thrusting. The corridor lies at the northern end of the Kingman Uplift, a Laramide-age northeast-facing topographic high. The Frenchman-Overton corridor contains an overprint of late Cenozoic tectonism that is so strong that the true location and geometry of those early boundaries can only be understood by reconstruction. Stratigraphically, the Frenchman-Overton corridor embraces a contact zone between thick Paleozoic, Mesozoic, and Cenozoic sedimentary rocks in the north and Precambrian crystalline basement rocks overlain by late Tertiary volcanic rocks forming the Colorado River Extensional Corridor in the south. As an important part of this contact zone, the corridor also encompasses the area of intersection between two major late Cenozoic strike-slip fault systems: the right-lateral Las Vegas Valley shear zone and the left-lateral Lake Mead fault system (Lake Mead fault system; Fig. 1B). We refer to the area of intersection as the Callville interaction zone, an area of extremely heterogeneous late Cenozoic deformation. Displacements on the strike-slip fault systems occurred during regional extension and are some of the largest and best constrained in the Basin and Range. One estimate places concurrent extension along the corridor at 60–80 km in a 250°–270° direction between ca. 16 and ca. 10 Ma (Duebendorfer et al., 1998), while a more recent interpretation is that the Frenchman Mountain block restores to the top of Gold Butte and therefore total extension of Frenchman Mountain was ~70 km to ~260° (Fryxell and Duebendorfer, 2005). The displacements were accompanied by Miocene and younger basin sedimentation on which this field trip is focused. The basin-fill sequences are deformed into structures that are dominated by normal and strike-slip faults, but also locally include common open to isoclinal folds of varied orientations and reverse faults. Although the late Tertiary sedimentary rocks are the main focus of the trip, the journey will provide numerous views into the northern extreme of the southern igneous terrain. STRATIGRAPHY OF THE FRENCHMAN-OVERTON CORRIDOR The stratigraphy of the Frenchman-Overton corridor (Fig. 5) has been described in detail by Longwell (1928, 1936), Longwell et al. (1965), and Bohannon (1984). The pre-Cenozoic strata are characterized by the dominantly marine carbonate Paleozoic cratonal rocks of the Grand Canyon region and their lateral miogeoclinal equivalents and the mostly non-marine, siliciclastic Mesozoic sequence. Tertiary strata typically lie in disconformable or in slightly angular unconformable contact with the underlying Paleozoic and Mesozoic units. Buttress unconformities are locally common. The Tertiary stratigraphy was thoroughly described by Bohannon (1984) and modified by Beard (1996). Although lithologically similar units have been classified and mapped throughout the area, Beard (1996) has shown that facies changes can

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occur over relatively short lateral distances, calling into question the lateral continuity and contemporaneity of the Tertiary strata in the Frenchman-Overton corridor. We will make use of the generally accepted late Tertiary stratigraphic nomenclature, but it is one goal of this field trip to highlight some of its shortcomings. The Horse Spring Formation The oldest and thickest Tertiary unit is the Horse Spring Formation (ca. 24–12 Ma) (Longwell et al., 1965; Bohannon, 1984; Beard, 1996; Figs. 5 and 6). It is a heterogeneous package of conglomerates, sandstones, mudstones, limestones, and evaporites (mainly gypsum) and has been subdivided into four members, all of which crop out in the Frenchman-Overton corridor. The Rainbow Gardens Member The Rainbow Gardens Member of the Horse Spring Formation crops out throughout the Frenchman-Overton corridor, from its type locality in the Rainbow Gardens Recreation Area southeast of Frenchman Mountain to the Virgin Mountains on the east (Figs. 5–7). Fission-track, K/Ar, and 40Ar/39Ar geochronology brackets the age of the Rainbow Gardens Member between ca. 26.0 and <18.8 Ma in the south Virgin Mountains (Beard, 1996). Bohannon (1984) reported that few dates were available for the Rainbow Gardens Member but placed its younger age at ca. 17.2 Ma based on a date from the base of the Thumb Member. This particular sample was reanalyzed and found to be 15.6 Ma, so the exact upper limit of the Rainbow Gardens Member is not well-constrained throughout the corridor. The vertical succession of the Rainbow Gardens Member is similar throughout its outcrop belt and consists of (1) a basal, clast-supported conglomerate with a dominantly carbonate clast composition, suggesting that it was sourced by the Paleozoic miogeoclinal sequence; (2) a flaggy-weathering, reddened, mudstone and sandstone sequence with planar laminated and cross-stratified sandstone beds; and (3) a resistant limestone and dolomite unit that principally is comprised of massive beds at its base and somewhat thinner beds of algal laminite near its top. In the Virgin Mountains area to the east and the northern Muddy Mountains to the north, the middle section includes volcaniclastic sandstones and ashfall tuffs. In the Rainbow Gardens Recreation Area, the basal conglomerate unconformably overlies the Upper Red Member of the Moenkopi Formation (Triassic), whereas farther east it may overlie the Permian Kaibab, Triassic Chinle, Moenave-Kayenta, Moenkopi or Jurassic Aztec Formations. This relationship suggests that the basal Rainbow Gardens Member contact is a throughgoing, major angular unconformity (Bohannon, 1984). In the Virgin Mountains area, Beard (1996) has interpreted a disconformity to angular unconformity that defines the top of the Rainbow Gardens Member as well. She interpreted pedogenic features in the lower portion of the carbonate sequence and local thin conglomeratic sandstones in the middle part of the carbonate sequence as indicative of a period of subaerial exposure. In this interpretation, the more massive, reddish-pink lower carbonate

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Development of Miocene faults and basins in the Lake Mead region Q

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Chinle Formation

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TERTIARY

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? Moderately to steeply dipping normal fault showing dip Ball on downthrown side; arrow near dip symbol shows trend of striae on fault surface .

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Anticline

°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°°° Tufa with silica in the Lovell Wash Member of the Horse Spring Fm. 52

Strike and dip of inclined bedding Horizontal bedding

15

Minor anticline showing bearing and plunge of hingeline Overturned syncline

Figure 6 (on this and previous page). Geologic map of the Lava Butte area, east of Frenchman mountain, after Castor et al. (2000). Numbered stars indicate location of Stops 1.2–1.5. See Figure 2 for location of Stops 1.1 and 1.6.

sequence is the uppermost portion of the Rainbow Gardens Member. An abrupt shift to algal laminite deposition and a lack of pedogenic features suggests that the upper carbonate sequence may actually be the basal unit of the Thumb Member of the Horse Spring Formation. This upper carbonate sequence exhibits abrupt facies changes laterally into more typical sandstone, mudstone, and gypsum facies of the Thumb. Beard (1996) interpreted the disconformity to be caused by the onset of extension at ca. 16 Ma. This hypothesis is yet to be tested fully in the FrenchmanOverton corridor. The Thumb Member The Thumb Member also crops out throughout the Frenchman-Overton corridor and ranges in age from ca. 16 to ca. 13.5 Ma (Bohannon, 1984; Beard, 1996; Figs. 5–10). In the south Virgin Mountains, this range is more tightly constrained between ca. 16.2–14.2 Ma (Beard, 1996). New 40Ar/39Ar dates (Donatelle et al., 2005; Martin, 2005) from the Thumb Member

within the Echo Wash area (Fig. 8) range from 16.36 (biotite) to 14.58 Ma (sanidine). Bohannon (1984) and Beard (1996) have shown that the Thumb Member of the Horse Spring Formation is an extremely heterogeneous unit, showing abrupt lateral and vertical facies changes. The unit also shows great variation in thickness from a few hundred m to 1300 m in the Echo Wash area (Bohannon, 1984). In the Virgin Mountains area, Beard (1996) demonstrated that the abrupt facies changes were the result of deposition during active segmentation of the landscape during extension. Lithofacies within the Thumb include gypsum, gypsiferous limestone and mudstone, fine-grained siliciclastics, sandstones and pebbly conglomerates, cobble to boulder conglomerates with Proterozoic, Cambrian, and Paleozoic (undifferentiated) provenance, and carbonate units. In contrast to the Rainbow Gardens Member, the Thumb Member lacks a typical vertical sequence. We will spend time on this field trip demonstrating this heterogeneity and exploring its implications.

M. Lamb et al.

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A

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Figure 7. Photographs of key outcrop features from Day 1 field stops in the Rainbow Gardens Recreation Area. (A) View looking north at the Rainbow Gardens and Thumb Members of the Horse Spring Formation at Stop 1.2. (B) Same as A, with interpretation. Thrc—Rainbow Gardens Member conglomerate submember; Thrr—Rainbow Gardens Red submember; Thrl—Rainbow Gardens Limestone submember; Thtl—Thumb limestone. (C) View southwest of the gypsum quarry and conglomerate facies within the Thumb Member at Stop 1.4. Thtc—Thumb conglomerate. (D) Close-up view of Thumb Member gypsum units shown in C; note the massive, bedded gypsum on the left that shows large-scale, bladed gypsum fabrics and more mud-rich gypsum facies lying stratigraphically below on the right. (E) Thumb Member conglomerate showing both matrix- and clast-supported textures, suggestive of debris flow and traction current sedimentation, respectively.

Development of Miocene faults and basins in the Lake Mead region The Bitter Ridge Limestone Member Small, jumbled and deformed outcrops of the Bitter Ridge Limestone Member have been mapped in the Rainbow Gardens Recreation Area, but this unit is best exposed in the Bitter Ridge area (west of the Longwell Ridges), its type locality (Figs. 5, 6, 8–10). Here, it is a thick-bedded, algally laminated limestone with teepee structures, oncolitic textures, and stromatolitic bioherms. Bohannon (1984) placed the age of the Bitter Ridge Limestone at between ca. 13.5 and 13.0 Ma, based on fission track ages. Castor et al. (2000) reported a 40Ar/39Ar age of 13.07 ± 0.08 Ma for an ash in the Bitter Ridge Limestone obtained south of Lava Butte in the Rainbow Gardens Recreation Area, but this age is younger than ages they obtained for the “overlying” Lovell Wash Member. If the small exposures in the Rainbow Gardens Recreation Area are ignored, most of the Bitter Ridge

The Lovell Wash Member The Lovell Wash Member crops out mainly in the Lovell Wash and White Basin areas; it has also been mapped in the Rainbow Gardens Recreation Area (Figs. 5, 6, and 8). Bohannon (1984) considered the Lovell Wash to range from 13.0 to

B as in F aul t

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fault

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Figure 8. Geologic map of the Echo Wash–Longwell Ridges area, after Beard et al. (2006). Also includes recent 40 Ar/39Ar ages (Donatelle et al., 2005; Martin, 2005). Numbered stars indicate location of stops for Day 2.

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EXPLANATION

M. Lamb et al. Central

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Figure 9. Correlation diagram of Thumb and Bitter Ridge Limestone Members measured sections for Stops 2.1, 2.2, and 2.3. Inset map is a highly simplified version of Figure 8 to show relative locations of measured sections. Also note that lithofacies are highly schematized. Sections 1 through 4 are approximately hung from a set of ashes, but ongoing major and trace element geochemical analyses of these and other ashes will provide a more solid datum in the future.

Development of Miocene faults and basins in the Lake Mead region 11.9 Ma. Çakir et al. (1998; M. Çakir, 2005, personal commun.) bracket its age between 13.3 and 12.1 Ma. Harlan et al. (1998) reported 13.28 ± 0.09 and 13.17 ± 0.10 from basalts interbedded with the Lovell Wash in Boulder basin and suggested it could be at least in part time correlative with the Bitter Ridge Limestone. Castor et al. (2000) also report multiple 40Ar/39Ar ages for tuffs in the Lovell Wash exposed on the east side of Frenchman Mountain that range from 13.40 ± 0.05 to 13.12 ± 0.24 Ma and indicate they could find no discernable difference in age between the Bitter Ridge and Lovell Wash in the Frenchman Mountain area. Tuffaceous sandstones and mudstones comprise the dominant Lovell Wash lithofacies. Red Sandstone Unit The red sandstone unit was named informally by Bohannon (1984) to include clastic sequences in White Basin and the Frenchman Mountain area that are unconformable with both the underlying Horse Spring Formation and the overlying Muddy Creek Formation (Figs. 5 and 6). Duebendorfer and Wallin (1991) expanded the range of the unit to include scattered exposures between Frenchman Mountain and White Basin and a mostly coarse clastic basin-filling sequence in Boulder basin south of the Las Vegas Valley shear zone. In general the unit is synvolcanic and syntectonic, filling small basins formed during extension from ca. 12–8.5 Ma and interfingering in its upper part with the volcanic rocks of Callville Mesa (described below). More recent 40Ar/39Ar work by Harlan et al. (1998) yielded an age of 11.70 ± 0.08 Ma from a tuff near the base of the unit in the Gale Hills, north of the Las Vegas Valley shear zone. The red sandstone is characterized by sandstone, interbedded with siltstone and claystone, is locally gypsiferous, and contains abundant thin white tuff beds. Conglomeratic facies and locally megabreccia are common near some basin-margin faults (Bohannon, 1984). Adjacent to basin margin faults such as the Muddy Peak fault on the west side of White Basin, the unit is conglomeratic with intercalated megabreccia deposits. Muddy Creek Formation The name “Muddy Creek Formation” has been applied widely to late Miocene deposits that are mostly flat lying and post-date extension (Figs. 5 and 6). Bohannon (1984) restricted the name to rocks that are demonstrably continuous with those at the type locality north of Glendale, Nevada. However, subsequent studies have continued to use the term in the Lake Mead area for separate basins of similar age (e.g., Castor et al., 2000, Duebendorfer, 2003; Anderson, 2003). The restricted Muddy Creek basin extends from about Echo Bay in the Overton Arm of Lake Mead northward into the Virgin Valley and Moapa area; its maximum extent represents several small basins that coalesced into one large basin by ca. 6 Ma or younger (Pederson et al., 2000). Deposits include fine-grained clastic rocks,

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interbedded calcareous mudstone, gypsum, and local minor conglomeratic facies at the basin margins. Pebble to cobble fluvial gravels are interbedded with the very top of the Muddy Creek at the tip of Mormon Mesa in Overton and probably reflect initial dissection of the Muddy Creek basin related to integration of the Virgin and Colorado Rivers by ca. 5 Ma. The Muddy Creek Formation exposed in the FrenchmanOverton corridor was deposited in small separate basins. In the Frenchman Mountain area, Castor et al. (2000) mapped fairly widespread facies of limestone, gypsite and gypsiferous gravel, fine-grained sandstone and siltstone, and pebble to boulder conglomerate. The deposits fine and thicken northward into Nellis basin, a depocenter north of Frenchman Mountain (Castor et al., 2000). Eastward in the Government Wash area fine-grained Muddy Creek deposits give way to cobble to boulder conglomerate deposits (Duebendorfer, 2003) that probably represent alluvial fan deposition. Further east in Boulder basin, Anderson (2003) mapped five separate depocenters containing as much as a few hundred meters of coarse conglomerate, sandstone, and locally gypsum and gypsiferous mudstone. These basins record fanning upward dip patterns, folds and unconformities, all indicating syndepositional deformation. We will visit one of these depocenters on the field trip. Hamblin-Cleopatra Volcano The Hamblin-Cleopatra stratovolcano was cut into three lobes and offset by as much as 20 km in a left-lateral sense by the Hamblin Bay strand of the Lake Mead fault system (Anderson, 1973; Figs. 3 and 11). Hamblin Mountain is the eroded western lobe of the original volcano on the northwest side of the fault, whereas the Cleopatra rocks, cut into two lobes, form the stranded eastern part. The Hamblin volcanics range in age from 10.07 to 11.7 Ma (40Ar/39Ar; Anderson et al., 1994) whereas ages from the Cleopatra lobe are 13.1–12.5 Ma (K/Ar; Thompson, 1985, M. Kunk, 2000, personal commun.). Because of differing dating techniques and unsystematic sampling, these differences are not considered significant and the volcano is considered to be 12–10 Ma. The rocks are mostly massive to flow-banded andesite flows and thick, lighter colored autobrecciated flows, cut by synvolcanic dikes, sills, and small plugs of porphyritic to equigranular andesite and dacite. Basal flows of the Hamblin lobe are interstratifed with strata of Lovell Wash Member (Anderson, 2003). The more intact easternmost Cleopatra lobe (Fig. 3) exhibits an extensive set of radial dikes that in some places comprise as much as 80% of the outcrops. Volcanic Rocks of Callville Mesa The Callville Mesa volcanic field, just west of Hamblin Mountain, includes basaltic andesite flows erupted from cinder cones, primarily in the Callville Mesa area in Boulder basin. The flows are interbedded with and overlie the top of the red

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D Figure 10 (on this and following page). Photographs of key outcrop features from Day 2 field trip stops in the Echo Wash–Longwell Ridges Area. (A) View east of repeated fault blocks of Thumb units in the Echo Wash area, southeast of East Longwell Ridge. (B) Same as A, with interpretation. Line of section is location for Stop 2.1 and measured section 1 on Figure 9. Ja—Jurassic Aztec formation; Thtg—Thumb gypsum; Thtl—Thumb limestone; Thtc—Thumb conglomerate. Note that these abbreviations are the same for units in the Rainbow Gardens Recreation Area but are not the same submembers nor are they even the same age. (C) View east of the Thumb conglomerate facies in depositional contact with Paleozoic carbonates forming a buttress unconformity at Stop 2.3 on the southern tip of West Longwell Ridge. (D) Same as C, with interpretation. Pz—Paleozoic; cg—conglomerate; ss—sandstone. (E) View west from Stop 2.3 of the Bitter Ridge Limestone Member overlying the Thumb Member of the Horse Spring Formation. (F) Same as E, with interpretation. This plate includes the location of the westernmost measured section 8 on Figure 9. (G) Close-up view of teepee structure in the Bitter Ridge Limestone Member at Stop 2.4. Scale bar is ~15 cm. (H) View northeast of the striations on the White Basin fault at Stop 2.5 with overlying Thumb beds in the hanging wall; note person for scale in left-central portion of photo. Trend and plunge of striations is 250°, 46°. (I) View northeast of upper Thumb facies of Stop 2.7 in Echo Wash; the light-colored beds on the right comprise a thick ash that shows a 40Ar/39Ar age of 14.85 ± 0.07 Ma. The interbedded fine-grained sandstones and mudstones on the left are characteristic of the Thumb sandstone facies and suggest episodic deposition in a subaerial (distal alluvial fan or marginal lacustrine?) environment.

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sandstone unit; lowermost flows are variably tilted, whereas the youngest flows are not. Feuerbach et al. (1991) interpreted the Callville field as erupting during the late stages of extension that formed the red sandstone depocenters of Boulder basin. Anderson (2003) suggested the Callville Mesa eruptive centers were satellitic to the Hamblin-Cleopatra volcano. Ages of the volcanic rocks of Callville Mesa range from 11.41 ± 0.14 Ma (40Ar/39Ar, Harlan et al., 1998) to 8.49 ± 0.20 (K/Ar, Feuerbach et al., 1991). The older age is from a basaltic andesite flow that overlies the red sandstone unit with a 20° to

40° angular discordance. The younger age is from the uppermost flow on Callville Mesa. MAJOR STRUCTURES AND TIMING OF DEFORMATION ALONG THE FRENCHMANOVERTON CORRIDOR We begin our transect on the eastside of Frenchman Mountain (Figs. 3, 4, and 6). The Frenchman Mountain block is an east-dipping, fault-bounded homocline of crystalline basement

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Figure 11. Tectonic map of the Callville Bay Quadrangle showing stops for Day 3 (numbered stars); faults bounding major structural blocks (heavy lines); fold axes (light lines); and localities where reverse faults or outcrop-scale isoclinal folds are exposed (dots).

overlain by Paleozoic to lower Mesozoic strata that are, in turn, overlain in low-angle unconformity by the Miocene strata on which the beginning of this trip is focused. Directly east of Las Vegas, the block is bounded by a large-displacement down-tothe-west normal fault called the Frenchman fault. Southward, the Frenchman fault curves to the southeast and becomes a right-normal fault. Northward, it curves to the northeast and interacts with the large-displacement Las Vegas Valley shear zone. Throughout its length, moderate to deep basins, hidden beneath the alluvium of Las Vegas Valley, have formed in the hanging wall of the Las Vegas Valley shear zone as indicated by geophysical data (Langenheim et al., 2001). At the east margin of Frenchman Mountain, the uniformly east-tilted Miocene sedimentary rocks are repeated by a normal fault and pass eastward, across north-northeast–striking faults, into Miocene rocks of the Boulder basin (Duebendorfer and Wallin, 1991; Castor et al., 2000). The Boulder basin and its fill are interpreted to have been formed as a result of movement along the kinematically linked Las Vegas Valley shear zone and Saddle Island detachment fault (Duebendorfer and Wallin, 1991). The

Saddle Island detachment fault is interpreted to be a major extensional structure in the Frenchman-Overton corridor. It is characterized by a Proterozoic, mylonitic, lower plate, and a brittlely deformed Miocene upper plate (Duebendorfer et al., 1998). The rocks of the Boulder basin south of the Las Vegas Valley shear zone and the Gale Hills north of the Las Vegas Valley shear zone (Fig. 3) are intensely deformed on a heterogeneous array of structures most of which are interpreted as reflecting strain accommodation associated with the interaction between the Las Vegas Valley shear zone and the Lake Mead fault system (Anderson et al., 1994; Duebendorfer and Simpson, 1994). One of the largest structures reflecting this interaction is a broad zone of faults and folds covering most of the Gale Hills that experienced clockwise steep-axis rotation (Sonder et al., 1994). Whether this strain reflects right-sense drag on the Las Vegas Valley shear zone or north-south contractional collapse, or a combination of both, is controversial (see additional discussion below; Anderson et al., 1994; Duebendorfer and Simpson, 1994; Çakir et al., 1998). Other large-scale structures include northerly striking strike-slip faults and east-west–trending folds, the largest of which is the Lovell Wash syncline. These structures do reflect north-south contractional strain. The last day of this trip is focused on the east extreme of the Boulder basin–Gale Hills areas and includes a visit to deformed roundstone gravels near Callville Bay suggestive of a continuation of deformation into the neotectonic regime. East of the Boulder basin–Gale Hills area extending to and including the Overton Arm basin, the geology is dominated by Miocene basins, the dismembered parts of a Miocene stratovolcano (Callville Mesa and Hamblin volcanics discussed above), and intervening blocks of pre-Tertiary rock distributed along strands of the Lake Mead fault system. The Lake Mead fault system comprises several sinistral-slip faults, which strike westsouthwest and run from Callville Bay to Overton Arm of Lake Mead and then northeast to the south Virgin Mountains (Fig. 3). This system includes the Bitter Spring Valley, Hamblin Bay, and Rodgers Spring faults west of the Overton Arm of Lake Mead (Fig. 3). The Overton Arm basin runs from the Bitter Spring Valley to the south Virgin Mountains (Fig. 3) and has been interpreted as a pull-apart basin between left-stepping left-lateral faults (Campagna and Aydin, 1994). The White basin is a major extensional basin north of the western end of the Overton Arm basin. Steepaxis clockwise rotations of some blocks in the western part of this area are indicated by unpublished paleomagnetic data. Much of the deformation associated with the structures in the Lake Mead area is thought to occur after the initiation of deformation to the south (ca. 27–15 Ma) and north (20–10 Ma; Fig. 1B; Spencer and Reynolds; 1989; Axen et al., 1990, 1993; Duebendorfer et al., 1998). The main deformation in the Lake Mead region occurred from 16 to 10 Ma (Anderson et al., 1972; Bohannon, 1984; Wernicke et al., 1988; Duebendorfer and Simpson, 1994; Beard, 1996; Fig. 8). Most deformation moved west at ca. 10 Ma to the Death Valley to Owens Valley part of the Basin and Range at this latitude. However, despite many of the structures in the Frenchman-Overton corridor providing

Development of Miocene faults and basins in the Lake Mead region good evidence for deformation diminishing at 10 Ma, locally significant deformation continued well after 10 Ma. For example, deformation in the Callville interaction zone continues at least through Muddy Creek time (10–6 Ma; Anderson, 2003; Duebendorfer, 2003) and may continue after 6 Ma (Anderson, 2003). In addition, low rates of deformation continue into the Quaternary in the Lake Mead region. INITIAL STUDIES AND CONCLUSIONS Initial geologic mapping in this region was completed by Longwell (1928, 1936), Longwell et al. (1965), and others in the 1960s. In his pioneering study in the Eldorado Mountains south of Lake Mead, Anderson (1971) recognized that significant extension resulted in steeply tilted strata displaced by low-angle normal faults. On the basis of reconnaissance stratigraphic studies and K/Ar analyses, Anderson et al. (1972) established that the volcanic strata to the south and the sedimentary strata to the north in the Lake Mead region were contemporaneous during the Miocene. This led to the understanding that (1) the contrasting strain between these two lithologically contrasting zones is also contemporaneous, (2) a major northeast-striking left-slip fault system (Lake Mead fault system) with lateral displacement of at least 65 km separates the central part of the differentially deformed terranes, and (3) the contrasting strains need to be integrated into a single tectonic model (Anderson, 1973). One tectonic explanation is that “large lateral displacements and extreme structural complexity record synextensional rafting of structural blocks atop a flowing undermass and concomitant contraction of the zone of flowage” (Anderson et al., 1994, p. 1381). In this model, the variety of types, scales, and attitudes of structures can be explained by this relatively long-lived process. Blocks to the north of Lake Mead have moved south, creating major shortening and the occlusion (tectonic escape) of material in an east-west direction in the Lake Mead area (Anderson and Barnhard, 1993). Zones of major translation and occlusion—such as the one in the Lake Mead area that resulted in juxtaposing coeval igneous and sedimentary domains—commonly have a low potential for preserving structural history because only the last rocks to arrive at any point are preserved. Sediments contained in adjacent syntectonic basin-fill sedimentary rocks can provide reliable measures of offsets or provide the only record of the former presence of vacated rocks. Much of the understanding of large translations in the Lake Mead area was either gleaned from or supported by such deposits (Anderson, 1973; Longwell, 1974; Rowland et al., 1990; Duebendorfer et al., 1998). In the western Lake Mead area, basin deposits retain a partial record of the original north-south transitional assemblage between the igneous and sedimentary terranes. Geologic mapping and structural studies south of Lake Mead (Anderson, 1971, 1977b, 1978b) together with chemical analysis of a detailed stratigraphic section of volcanic rocks (Anderson, 1977a, 1978a) led to the understanding that the onset of major deformation was not accompanied by significant shifts

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in major-element chemical composition of volcanic rocks. North of Lake Mead, by contrast, comprehensive lithostratigraphic studies (Bohannon, 1979, 1983b, 1984; Beard, 1996), led to the understanding that the Miocene sedimentary rocks record a transition from pre-extension deposition in a broad sag basin to synextensional deposition within more restricted basins developed in a mixed strike-slip and normal-faulting setting (Beard, 1996). This sag basin was dissected by normal and strike-slip faulting at the beginning of deposition of the Thumb Member (ca. 15.7–16 Ma; Fig. 5) into numerous subbasins. She called on “mixed mode” (normal and strike-slip) faulting to explain the complex facies variations that she detected in the Miocene Horse Spring Formation. Subsequent reconstructions suggest that 60–80 km of extension from Miocene faulting has occurred (Duebendorfer et al., 1998; Fryxell and Duebendorfer, 2005), implying that the Frenchman Mountain block likely lay either immediately west of the south Virgin Mountains or on top of Gold Butte when extension began. Therefore, the sag basin was much smaller than any configuration based on its present outcrop distribution. RECENT STUDIES AND CONTROVERSY The first several decades of studies of the excellent exposures of rocks and heterogeneous structures in the Lake Mead region created a fertile breeding ground for tectonic models. Controversy abounds! Many issues remain unresolved. The relative importance of normal versus strike-slip faults is debated (e.g., Anderson, 1973; Ron et al., 1986; Duebendorfer and Black, 1992; Duebendorfer and Simpson, 1994; Wawrzyniec et al., 2001), as are the details of how and why faults develop through time throughout the entire area and, in particular, within the Callville interaction zone and the Boulder basin just south of the Las Vegas Valley shear zone (Anderson, 1973; Bohannon, 1979; Duebendorfer and Wallin, 1991; Duebendorfer and Black, 1992; Anderson et al., 1994; Duebendorfer and Simpson, 1994; Çakir et al., 1998). The tectonic escape model, discussed above, explains the development of a variety of structures through time with one continual process. A different tectonic model by Duebendorfer and Simpson (1994) emphasizes the development of different sets of structures through time. The late-stage (ca. 12–8 Ma) sedimentation in the Boulder basin is interpreted to reflect large-magnitude extension on a regional-scale detachment fault mapped at Saddle Island in the southwest extreme of Lake Mead (Duebendorfer and Black, 1992) and to be kinematically linked to right-sense displacement on the Las Vegas Valley shear zone (Duebendorfer and Simpson, 1994). Duebendorfer and Simpson (1994) recognized the development of contractional strain but concluded that it developed after the major extensional features. Additional work since 1994 has not resolved this debate. For example, recent geophysical studies west of Frenchman Mountain (Langenheim et al., 2001) failed to reveal evidence either for a regional detachment fault or for reported large-magnitude extension beneath Las Vegas Valley (Wernicke et al., 1988). Recent detailed mapping of the east and west extremes of Boulder basin

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failed to reveal structures consistent with large extension above a detachment fault (Anderson, 2003; Beard et al., 2006). In addition, basinal settings for the Miocene strata are complex and not well understood. Much of the current research is focused on trying to better understand this process of basin dismemberment and development in a system of linked strike-slip and normal faults. Past models have suggested pull-apart basins (Duebendorfer and Wallin, 1991; Campagna and Aydin, 1994), rift-style extensional basins (Bohannon, 1984; Beard, 1996), and local contractional or transpressional basins (Anderson, 2003). The pull-apart basin suggested by Duebendorfer and Wallin (1991) applied to the ca. 12–8 Ma red sandstone in Boulder basin (more accurately described as a basin formed by the kinematically coupled motion between a strike-slip and detachment fault [E.M. Duebendorfer, 2005, personal commun.]). The pull-apart basin suggested by Campagna and Aydin (1994) applied to the Overton Arm basin; these authors did not give a time for the formation or development of this basin, but the argument that present gravity data suggest a basin indicates that it must have formed late in the regional deformation. As noted above, Beard (1996) emphasized mixed strike-slip and normal fault-related sedimentation for the Thumb Member in the south Virgin Mountains, but this hypothesis has not yet been tested throughout the region. Our current work in the Longwell Ridges area suggests that the onset of extension during late Rainbow Gardens or early Thumb time produced small, rapidly subsiding basins with abrupt lateral facies changes. We are also documenting sets of cross-cutting faults throughout the basins that record a geological complexity and history that appears similar to, and should help inform our understanding of, the Callville interaction zone. As these basins are mapped, dated, and delineated in detail, a better understanding of the structural and tectonic evolution of the Frenchman-Overton corridor and Lake Mead region will be gained. The following questions delineate many of the unresolved structural and stratigraphic basin-related issues: 1. Does Boulder basin in the western Lake Mead area represent major (Duebendorfer and Wallin, 1991) or minor extension (Anderson and Barnhard, 1993)? If the basin represents major extension, where and how does that extension die out to the south? Is the basin a pull-apart basin? If it is, for what part of the stratigraphic section does that model apply? If it is not, what type of basin was it? 2. Is contractile shortening a local phenomenon restricted to a narrow (~300 m) corridor along the Las Vegas Valley shear zone (Duebendorfer and Simpson, 1994); i.e., the “Callville Bay interaction zone,” or is it a province-scale phenomenon (Wernicke et al., 1988, Anderson and Barnhard, 1993, Çakir et al., 1998; Anderson, 2003)? 3. Do the major strike-slip faults (the Las Vegas Valley shear zone and Lake Mead fault system) transfer strain from adjacent or along-strike areas of contrasting amounts of extension (Liggett and Childs, 1977; Duebendorfer and Black, 1992; Duebendorfer and Simpson, 1994; Duebendorfer et al., 1998)? Or are they first-order structures that

control the distribution and nature of extension (Ron et al., 1986; Campagna and Aydin, 1994)? Or are they the boundaries of a west-widening tectonically escaped block (Anderson and Barnhard, 1993)? If there is an escaping block, what are its boundaries? Does the Las Vegas Valley shear zone have an earlier, pre-Miocene history? 4. Was the long-distance (ca. 65–70 km) westerly displacement of Frenchman Mountain accomplished mainly by detachment faulting (upper-plate active) (Duebendorfer and Wallin, 1991; Duebendorfer and Simpson, 1994) or was it rafted passively on a westerly directed current of ductilely flowing undermass (Anderson, 1973; Anderson and Barnhard, 1993)? Or are there new models that better explain the displacement? 5. What are the ages, depositional environments, and structural settings and configurations of Miocene sedimentary sequences and basins throughout the region? Can a better-constrained regional stratigraphy be defined? Do basins record an east to west progression of extension as some models would predict, or was basin initiation synchronous across the region? Do basins record major times of fault (and basin) reorganization, or was the region one continually evolving system of strike-slip and normal faults? Does the history of basin development help us answer the structural-tectonic controversy outlined in (3) above? Do basins indicate that tectonic processes dominated the development of the region, or were climatic processes also important? Any tectonic model that hopes to explain the complex structural and stratigraphic relationships in the Lake Mead corridor must address the following: • The Lake Mead fault system and its associated faults with a left-lateral sense of offset. • The Las Vegas Valley shear zone and its apparent sense of right-lateral offset (although Anderson will point out on this trip that its most recent motion in the Lake Mead region may be left lateral!). • Folds that are common in the Callville interaction zone and that affect strata as young as the late Miocene to Pliocene Muddy Creek Formation (and possibly Pliocene Colorado River deposits) in the vicinity of Callville Bay. • Northeast-southwest–striking to north-south–striking normal faults that offset Miocene strata. • The contrast between Miocene sedimentary strata to the north and igneous units to the south of the Las Vegas Valley shear zone and Lake Mead fault system. • Chronostratigraphic and facies relationships of the Miocene sedimentary basin fill. Although these lists may seem daunting, we believe present and future basin analysis studies, coupled with study of the local faults, have a high potential of contributing to and answering these questions. Tectonic models to date have been developed based on geologic mapping and structural analysis at a range of scales, an absolutely necessary precursor to the development of accurate

Development of Miocene faults and basins in the Lake Mead region models and a clear understanding of the processes that accompany extension in this region. All data suggest that the early Miocene is a major transition time in terms of the tectonic evolution of the area. Metamorphic core complexes were “shutting off” to the south, major extension was ending to the north, and the strike-slip and normal faulting processes that characterize the Lake Mead corridor were “turning on.” Fortunately, Miocene sedimentary strata are beautifully exposed throughout this area and provide valuable clues to understanding its tectonic evolution. We suggest that the first stage of geologic mapping and the development of an early phase of tectonic models for this area are coming to a close. In their place we find the need for more detailed (1:24,000 and larger scale) mapping, high resolution stratigraphic correlation at the lithofacies scale, and a more dense database of geochronological control. Further geophysical and geochemical-isotopic data would also substantially contribute to this effort. This should allow us to uncover the Neogene paleogeography more completely, address the questions listed above, and as a result, develop better-constrained tectonic models of this important transition zone.

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Day 1: The Frenchman Mountain Block and the Rainbow Gardens and Thumb Members of the Tertiary Horse Spring Formation (Figs. 2 and 6)

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The goal of the first day is to examine the oldest Miocene units in the Lake Mead region and examine a few large, perhaps less controversial structures in the area. After a short overview stop, we will spend most of the morning walking through a beautifully exposed section of the Rainbow Gardens Member. We will then examine a newly recognized unconformity within the upper Rainbow Gardens at a different location, introduce a few of the facies within the Thumb Member, and stop to observe two of the faults bounding the Frenchman Mountain block.

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Directions Leave Fiesta Hotel and head east (right) on Lake Mead Parkway; great views of the Las Vegas valley and Frenchman Mountain on the left. Turn left onto Mohawk Drive and immediately right into the broad graded north shoulder of Lake Mead Parkway and park. Walk 100 m north across the wash to the top of a low constructed ridge.

Stop 1.1: Overview Stop Lake Mead Parkway—Day 1 Overview Purpose To provide an overlook of Las Vegas Valley, Frenchman Mountain, and River Mountains area and present an introduction to the trip in general and the first day in particular.

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Directions Continue driving east on Lake Mead Parkway. On the left is the Lake Las Vegas Parkway to Lake Las Vegas resort. Entrance to the Lake Mead National Recreation Area. View to west in middle ground of prominent black butte is Lava Butte, a laccolith that intrudes Miocene section. Turn left onto North Shore Rd. Driving through Muddy Creek Formation. Turn left onto Lake Mead Blvd and continue past the entrance station (now leaving the Lake Mead National Recreation Area and entering Bureau of Land Management land.) View of Quaternary gravel channels overlying Muddy Creek Formation. Turn left onto Sunrise management area gravel road. Frenchman Mountain is tallest peak to the west. In the foreground to the west, view the orange-red and white ridge of the Rainbow Gardens Limestone submember. View of modern channel that exposes Quaternary gravel over reddish Thumb Member sandstone and conglomerate tilted to the southeast. Turn right at T-junction of two gravel roads onto Rainbow Gardens road. Cross through a ridge of Rainbow Gardens Limestone. View to the east is Rainbow Gardens Red submember. View Rainbow Gardens Conglomerate offset ~100 m along a left-lateral oblique fault. Triassic Moenkopi Formation is stratigraphically below this to the west. Stop 1.2: view of Lava Butte to the east.

Stop 1.2: Rainbow Gardens Recreation Area—Typical Section of the Rainbow Gardens Member, Horse Spring Formation Purpose To investigate and familiarize participants with the Rainbow Gardens Member of the Horse Spring Formation and the lowermost portion of the Thumb Member. We will also take a close look at the unconformity at the base of the Tertiary. Comments At this stop we will walk a west-east transect that begins in the Triassic Moenkopi Formation and passes stratigraphically upward through the lowermost member of the Tertiary Horse Spring Formation, the Rainbow Gardens Member. We will focus on the three submembers of this unit. Immediately west of where

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we’ll park is a resistant, ridge-forming carbonate, the Virgin Limestone Member of the Moenkopi. This parking location is in the Schnabkaib Member of this same formation, a yellow and white carbonate and gypsum-rich unit. We will walk east, across the valley and up the prominent ridge. As we cross the valley, an abrupt shift occurs, as the whitish-yellow Schnabkaib units give way to the Upper Red Member of the Moenkopi, characterized by ripple cross-laminated, fine-grained sandstones and dark reddish-brown mudstones. The top of the Moenkopi is overlain unconformably by the Rainbow Gardens Member. The ridge to the east comprises the three submembers of the Rainbow Gardens Member (Fig. 7A and 7B). We will begin by examining the Rainbow Gardens Conglomerate. The Rainbow Gardens Conglomerate comprises the dark gray, cliff-forming unit at the base of the slope. This moderately well-sorted, clastsupported, 2–20-m-thick, cobble conglomerate is made up of primarily Paleozoic limestone clasts with occasional sandstone clasts, presumably Jurassic Aztec (Ja) in origin. The red sandstone matrix is most likely also derived from the Aztec or Moenkopi. For the most part, this unit is reasonably well layered and somewhat organized, suggesting a predominance of tractional transport; there is very little evidence for en masse deposition as debris flows. Bohannon (1984) interpreted this submember as an alluvial deposit of a “gravel veneer on a widespread pediment surface.” More specifically, the Rainbow Gardens Conglomerate most likely represents the deposits of braided streams or a coalescing series of alluvial fans (a bajada) that formed on an erosional surface. The Rainbow Gardens Conglomerate is overlain by the much thicker, middle Rainbow Gardens Red submember, a mixed bag of pink-weathering limestone, red calcareous sandstone, and mudstone (Fig. 7B). Beds range in thickness from a few decimeters to a meter and are somewhat laterally continuous. Bohannon (1984) suggested this unit represents lacustrine, playa, and alluvial deposition. The uppermost submember, the Rainbow Gardens Limestone, holds up the ridge where we are headed (Fig. 7B). At this spot, the limestone is yellow-weathering. Just to our west, there is another prominent ridge of Rainbow Gardens Limestone that is pink-weathering with abundant stromatolites. Bohannon (1984) lumped the entire limestone ridge as Rainbow Gardens in age. In the Virgin Mountains, Beard (1996), however, documented an unconformity within this limestone section and placed the upper contact of the Rainbow Gardens Limestone with the overlying Thumb in the middle of this limestone unit. Hickson and Lamb have found evidence for a possible correlative unconformity in this locality, which we will examine here and at Stop 1.3. This new contact places the upper, stromatolitic limestone within the Thumb Member. The paleogeography during deposition of this member, as interpreted by Beard (1996), consisted of a regionally extensive and continuous sag basin. The basin held a large lake that may have expanded and contracted through time, producing the range of lacustrine, marginal, and fluvial facies present. This

is supported by the lateral continuity of the three submembers of the Rainbow Gardens throughout the Lake Mead region. In general, the same sequence of units visible at this stop can be correlated with nearly identical sections in the Virgin Mountains (Bohannon, 1984; Beard, 1996). While at this stop, we can discuss the following: What are the implications of strong lateral continuity of lithofacies? Is the Rainbow Gardens Conglomerate representative of a longitudinal (axial) drainage system or is it more of a classic bajada? Should the Rainbow Gardens be considered its own formation? Cumulative mi (km) 0.0 1.8

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Directions Turn around and drive north. Park on right just before a roadcut in the Rainbow Gardens Limestone.

Stop 1.3: Rainbow Gardens Recreation Area—Previously Unrecognized Unconformity in the Upper Part of the Rainbow Gardens Member Purpose To investigate in more detail a potential unconformity in the limestone submember of the Rainbow Gardens Limestone that could indicate an extended period of subaerial exposure at the top of this unit and to understand its implications. Comments We will make a short stop to see additional evidence for an unconformity within the Rainbow Gardens Limestone. Beard (1996) provided strong evidence for a previously unrecognized unconformity within the Rainbow Gardens Limestone that suggests that the lower Rainbow Gardens Limestone may be Rainbow Gardens age, whereas the upper Rainbow Gardens Limestone may actually be of Thumb age. Evidence for this unconformity in the Virgin Mountains includes a pedogenically altered limestone overlain by an algal-laminated limestone. Similar evidence is present at this locality, including rubification of the lower portion of the Rainbow Gardens Limestone, root traces and casts, possible mottling, and obliteration of primary sedimentary features. From up the hill on the north, we will look across the road to the south and view irregular surfaces within the Rainbow Gardens Limestone/Thumb above the possible unconformity that suggest paleotopography, possibly due to algal build-ups. Does the base of the Thumb Member actually lie at this unconformity? If this unconformity is truly regional in extent, as it appears to be, then what does this imply for tectonics? Cumulative mi (km) 0.0 2.9

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Directions Turn around and drive south. View Virgin Limestone Member of the Moenkopi Formation on both sides of the road. It is doubled here due to faulting.

Development of Miocene faults and basins in the Lake Mead region 4.4 6.7

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Leaving strike ridges. Turn left at the junction with the gravel road to Las Vegas. (This road is the southern extension of S. Hollywood Blvd. and can be used to return to Las Vegas.) “Reenter” Horse Spring ridges and faults Thumb conglomerate to the east and mustard-colored dolomite and sandstone of lower Thumb to the west. Stop 1.4.

Stop 1.4: Rainbow Gardens Recreation Area—Gypsum and Conglomerate in the Thumb Member of the Horse Spring Formation, Directly South of Lava Butte Purpose To examine the lower lithofacies of the Thumb Member of the Horse Spring Formation at its type locality. In particular, we will look at carbonate, gypsum, and conglomerate facies. Comments To the west at this stop is a mustard yellow unit that directly overlies the Rainbow Gardens Limestone and forms the prominent dipslope. This unit is the lowest exposed lithofacies of the Thumb Member of the Horse Spring Formation and consists of interbedded sandstones, mudstones, and carbonate beds. It appears to lie in gradational contact with the Rainbow Gardens Limestone in this locality. Immediately east is a large, north-south–trending pit: an abandoned gypsum mine (Fig. 7C). The working face lies to the south and can be clearly seen from the east side of the road. This face comprises massive, bedded gypsum, another of the key lithologies of the Thumb Member (Fig. 7D). We will walk north along the road to the roadcut exposure of the Thumb conglomerate, which crops out as a jagged, toothy ridge directly east of the gypsum pit. The Thumb conglomerate (Fig. 7E) here differs from the Rainbow Gardens Conglomerate that we investigated in Stop 1.2 principally in terms of its clast composition. Although still dominated by limestone and dolomite clasts derived from the Paleozoic carbonates, this conglomerate contains a subpopulation of igneous and metamorphic clasts that most likely were derived from units to the south. In general, however, this conglomerate is difficult to distinguish from the Rainbow Gardens Conglomerate units below and, as will be seen at the stops on Day 2, can actually be completely indistinguishable. If the unconformity recognized by Beard (1996) is the base of the Thumb Formation, then a vertical succession of the Thumb at this locality is as follows: • Algally laminated, creamy-white limestones and dolomites with occasional stromatolitic beds; • Interbedded sandstones, mudstones, and carbonates of the mustard-colored lithofacies; • Massive bedded gypsum; • Cobble conglomerate.

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This succession is not found at other Thumb localities, as we will see on Day 2 (Bohannon, 1984; Beard, 1996). The algal limestones most likely formed in a lacustrine setting. The overlying interbedded mixed siliciclastic-carbonate sequence suggests a transition to a more marginal lacustrine environment: the carbonates may represent background precipitation within the lake interrupted by periods of clastic input from fluvial sources. This sequence grades into the gypsum, which may represent a deepening of the lake. Although the depositional environment of the gypsum is not firmly established, the presence of large bladed gypsum crystals suggests quiescent deeper-water sulfate precipitation. Finally, the conglomerate represents a rapid shoaling at this locality. Perhaps the single most significant observation that can be made here is that the lithofacies of the Thumb that are exposed in the Rainbow Gardens area vary considerably laterally. This stands in sharp contrast to the Rainbow Gardens Member, with its relatively monotonous sequence of conglomerate, red sandstone and mudstone, and limestone that is found in all fault blocks containing the Rainbow Gardens Member. Is there really a “type section” of the Thumb, given its extreme lateral variability? Is it possible to even define one type section? Cumulative mi (km) 0.0

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Directions Continue driving northeast toward Lava Butte. Stop 1.5.

Stop 1.5: Rainbow Gardens Recreation Area—Major North-Striking Fault Purpose To view a northerly trending fault zone separating east-dipping sandstone and gravel of the Thumb Member on the west from steeply north-dipping younger Bitter Ridge Limestone beds on the east. Comments The sharp discordance between the Bitter Ridge beds and the fault results from left-sense drag of beds into the northerly striking fault. Left-sense slip is determined not only from the stratal bending seen here but also from fault fabrics and striations seen elsewhere. This fault is interpreted to have large displacement because it separates an extensive structural block to the east in which the Miocene sediments are interstratified with mafic lava flows, cut by dikes, and folded from the nonigneous and nonfolded Frenchman Mountain block to the west. If that interpretation is correct, this fault appears to be paired with the northwest-striking right-lateral portion of the Frenchman fault (observed from our introductory overview stop along Lake Mead Pkwy) to accomplish south-directed motion of the entire Frenchman Mountain block. Stated differently, the south-pointing prong of the Frenchman Mountain block appears to be inserted

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southward into an area where the basin sediment section contains igneous rocks. Additional basin studies, including paleomagnetic study, are required to test this hypothesis. Cumulative mi (km) 0.0

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Directions Turn around and drive south, retracing route to edge of outcrops; i.e., where the ridges abruptly end. Stop 1.6.

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Stop 1.6: Overview of the Frenchman Fault Purpose To provide a southward view of the steep (~80°) northeast dip of the Frenchman fault and discuss the overall geometry of this major fault. Comments The Frenchman fault has km-scale normal down-to-the-west throw directly east of downtown Las Vegas where, at the surface, it dips ~30° west. Geophysical data are consistent with a downward-steepening geometry of the fault in that area. At this stop, it is mainly a steep right-lateral fault. Within 1 km to the southeast, it dips as little as 65° northeast beneath the Frenchman Mountain block and is a right-reverse fault. We will have a brief discussion of the possible significance of this highly varied geometry and sense of slip and how they may relate to the fault viewed at Stop 1.5. Directions End of day. We will return to the North Shore Road and take this east to Echo Bay Resort (see Fig. 2). Day 2: Tertiary Horse Spring Formation Stratigraphy and Structure in the Vicinity of East and West Longwell Ridges (Fig. 9) The goal of the second day is to examine the Horse Spring Formation and observe important structures in the Longwell Ridges area, which is in the eastern part of the Frenchman-Overton corridor, or central part of the Lake Mead domain. The first two stops, in conjunction with the stops on Day 1, have been chosen to highlight the variability within the Thumb Member and issues related to assignment of local sections to the Rainbow Garden versus Thumb Members of the Horse Spring Formation. Additional stops will introduce the Bitter Ridge and Lovell Wash Members. One stop will feature a world-class exposure of the White Basin fault. Cumulative mi (km) 0.0 1.1

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Directions Leave Echo Bay Resort Looking west at East Longwell Ridge: dark Pennsylvanian-Permian Bird Spring Forma-

tion high to south overlying light-colored Mississippian Monte Cristo, which overlies the Mississippian-Devonian Sultan Formation. The Rogers Spring fault system bounds the front (east side) of the ridge and splays to the south into the various faults along the Longwell ridges. Turn left on to North Shore Rd. Echo Wash Take a right onto a dirt road into Echo Wash; we will be driving past the upper part of the Thumb Member for the next few miles. Reset odometer. Start of dirt road into Echo Wash. Go right ~0.4 mi up a side road off the main Echo Wash Road and park. We will walk a west-to-east transect into the Miocene section exposed at this locality. We will then walk along the lower part of the Miocene strata out to the south.

Stop 2.1: Echo Wash, Bitter Spring Quadrangle—Walk through Section of the Thumb Member Purpose To demonstrate the striking degree of lateral facies variations in the lower part of the Thumb(?) Member of the Horse Spring Formation and to show the time transgressive nature of the basal Miocene unconformity; to discuss whether this package of rocks should be assigned to the Rainbow Garden or Thumb Members; and to present new 40Ar/39Ar age data and discuss its implications for sedimentary basin evolution in the Tertiary. Comments At this locality (Figs. 9, 10A, and 10B), we will walk across the basal unconformity that places Tertiary strata atop Jurassic Aztec Sandstone. Conglomeratic, gypsiferous, and carbonate strata comprise the Miocene section (Fig. 9). We will begin by walking through a 100-m-thick section of Thumb in the area that Martin, Hickson, and Lamb have been working and mapping at a scale of 1:5,000. This section includes several informal subdivisions (in stratigraphic order as shown on column 1 on Fig. 10; note that on the map in Fig. 8 these units are lumped into two groups: the Thtc includes the lowermost conglomerate and sandstone and the Thlg includes the gypsum and limestone units): Ttc: conglomerate; Tts: sandstone; Ttrg: red gypsum; Ttwg: white gypsum; and Ttl: limestone. The basal conglomerate (Ttc), which lies in unconformable contact over the Jurassic Aztec Sandstone, shares all of the characteristics of the Rainbow Gardens Conglomerate examined at Stop 1.2. The ridgeline to the east is held up by a thin algallaminated limestone that shares many of the same features as the Rainbow Gardens Limestone in its type section. Furthermore, the thickness and overall physical stratigraphy of this sequence

Development of Miocene faults and basins in the Lake Mead region are somewhat comparable to the Rainbow Gardens Member in its type locality. As a result, this section was originally mapped by Bohannon (1983a, 1983b) as the Rainbow Gardens Member of the Horse Spring Formation. However, the intervening red and white gypsiferous facies differ strongly from the Rainbow Gardens, and the overlying carbonate units lack evidence for Beard’s (1996) unconformity (that now seems to be somewhat diagnostic of the Rainbow Gardens Limestone). The abrupt lateral changes in thickness of the overlying limestones, from ~10 m at section 1 to ~125 m at section 3 (Fig. 9; a 120% thickness increase over ~1 km distance!), is more diagnostic of the Thumb Member than the Rainbow Gardens Member. Beard (1996) shows thickness changes in the Rainbow Gardens carbonates, but they occur over much greater lateral distances. Finally, although there appears to be much lateral variability in the middle part of the Rainbow Gardens Member (Beard, 1996), it is predominantly a siliciclastic unit with minor carbonates and very minor gypsum; nowhere is the Rainbow Gardens Red submember predominantly evaporitic. Consequently, Beard et al. (2006) mapped the gypsum and carbonate units at this locality as the Thumb Member of the Horse Spring Formation, while assigning the conglomerate to the Rainbow Gardens Member. Our 40Ar/39Ar dates support Beard’s (1996) interpretation, with an age from the ash immediately overlying the sandstone interval above the conglomerate of 16.36 Ma and an additional date from within the red gypsum unit of 16.19 Ma. In addition, there is no significant evidence for a disconformity atop the conglomerate and, given our new ages from within the conglomerates elsewhere in this area, this supports a Thumb age for the conglomerate at this locality as well. However, nowhere is the upper Rainbow Gardens Member well dated; thus, the boundary between the Rainbow Gardens and Thumb Members is not well constrained. It can be argued, therefore, that the similarities in physical stratigraphy and generally poorly constrained ages for the Rainbow Gardens Member make the section at this stop a Rainbow Gardens candidate. This difference in interpretation highlights the need for a better-constrained chronostratigraphic framework. Facies are variable along strike (Fig. 9), and much of this appears to be due to abrupt lateral facies changes between the basal conglomerate and the overlying and laterally interfingering sandstone and minor limestone facies. This is particularly evident at this location as one examines the sandstone that overlies the conglomerate. The gypsum facies above us also changes laterally to the north into a sandstone-dominated facies, which then changes into conglomerate and sedimentary breccia at the base of the section. The overall stratigraphic pattern at this stop is one of a major buttress unconformity that that shows progressive onlap to the north. The 16.36 Ma (biotite) ash from this location provides a time horizon that can be traced across this section to the south and shows that the underlying conglomerate had marked topography on its upper surface. The overlying sandstone facies (in which the dated tuff occurs) interfingers to the north with the conglomerate; the sandstone thickens dramatically to the south in the stratigraphic interval below the ash. Paleotopography atop the

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conglomerate does not appear to be strongly erosional, given that individual conglomeratic stringers can be traced laterally into the sandstones, where they are subsequently deformed by soft sediment deformational processes. Cumulative mi (km) 0.0

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Directions Continue driving up (NW) the right branch of Echo Wash. The small track to right is the turn for Stop 2.1; continue past this, and stay right at all “Y” junctions. Stay right; Echo Wash turns northwest toward West Longwell Ridge (now directly ahead); the east end of Bitter Ridge is at 11 o’clock. The red sandstone of the Thumb Member is to the north. On the right at 3 o’clock is East Longwell Ridge, with folds in the Upper Paleozoic Bird Spring Formation, that was thrust over the Jurassic Aztec Sandstone in the Cretaceous. Beginning of the red sandstone of the Thumb Member on the left. Stop 2.2—view south of red sandstone with debris flow breccia atop a hill shaped like a “bunny”; this outcrop correlates with an outcrop directly north and with the red sandstone ridge to the northeast in the middle ground.

Stop 2.2: Bitter Spring Quadrangle—Brief Stop to Examine More Sand-Rich, Younger Thumb Member Facies (Optional) Purpose To show correlation of younger, coarse-grained sandstone and conglomerate facies, which are part of the upper Thumb Member. Comments We will stop briefly, time permitting, to point out several low ridges of interbedded conglomerate, sandstone, shale, and limestone. Within these ridges, several sections have been measured and correlated, in part to document facies transitions within the Thumb and in part to document faulting (Donatelle et al., 2004, 2005; Martin, 2005). The hill to the south of the road can be correlated with the sandstone and conglomerate ridge north of the road. These units have a distinctive suite of subangular conglomerate, matrix-supported coarse-grained sandstones, and crossstratified very coarse to granule size sandstones that allow us to correlate these units with a north-south–trending ridge that lies immediately east of West Longwell Ridge. Primary sedimentary structures and textures in these clastic units point toward deposition by sandy debris flows and moderately deep unidirectional tractional currents, implying deposition in a medial to distal

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alluvial fan environment. Recent 40Ar/39Ar ages (15.52 Ma and younger) indicate that these units are younger than all of the units investigated at Stop 2.1. Cumulative mi (km) 0.0 0.5

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Directions Continue driving up Echo Wash. Stop 2.3—stop just inside the southern mouth of the canyon between West Longwell Ridge and the Bitter Ridge, near the buttress unconformity.

Stop 2.3: Bitter Spring Quadrangle—Buttress Unconformity and the Bitter Ridge Limestone Member of the Horse Spring Formation Purpose To investigate the buttress unconformity of the Thumb conglomerate against the Paleozoic Bird Spring and Monte Cristo Formations; to examine (from afar) the Bitter Ridge Limestone Member of the Horse Spring Formation; and to present and discuss the implications of new geochronological data for these units. Comments This stop lies near the mouth of a canyon formed between Carboniferous carbonate units on the east and the Tertiary Bitter Ridge Limestone Member of the Horse Spring Formation on the west. The canyon roughly follows the trace of the White Basin fault. On the east side of the canyon the Pennsylvanian-Permian Bird Spring Limestone crops out as a thick-bedded, dark gray unit that dips to the southwest. In this locality it is petroliferous and contains a bryozoan indicator fossil called chaetetes. The Monte Cristo limestone lies stratigraphically below the Bird Spring and comprises the majority of the eastern canyon wall. Most striking at this locality is the impressive buttress unconformity exposed immediately to the east (Figs. 10C and 10D). Basal conglomerates of the Thumb Member onlap the dipping and eroded Paleozoic strata. Clast compositions of this conglomerate are consistent with a local, Paleozoic carbonate source, and grain size at this locality ranges from cobbles to large boulders, with the largest clasts located near the buttress contact. 40 Ar/39Ar dates (Donatelle et al., 2005; Martin, 2005; Fig. 8) from within the conglomerate at this locality (15.52 ± 0.07 Ma) and measured sections from the east side of this ridge (14.58 ± 0.06 and 14.91 ± 0.10 Ma) indicate these conglomerates and the overlying clastic sequence are younger than the Thumb units examined at Stop 2.1 (Fig. 8). Indeed, these rocks are correlative with the clastic sequence that overlies the Thumb Member limestone exposed in Echo Wash east of Stop 2.1. Mapping by Martin shows that this buttress unconformity continues northward along the east side of West Longwell Ridge, with progressively higher stratigraphic intervals lapping onto the Paleozoic unconformity. To our knowledge, the buttress

mainly places conglomerate against Paleozoic limestone, but the conglomerates abruptly give way laterally to sandstone and mudstone facies. To the west, the Bitter Ridge Limestone Member of the Horse Spring Formation comprises the spectacular cliff face. In this locality, the Bitter Ridge Limestone overlies a Thumb sequence of sandstones and mudstones (red units) and evaporites (whitish units) comprised mainly of gypsum (Fig. 10E and 10F). The basal portion of the Bitter Ridge Limestone is mainly a clastic sequence of thick- to medium-bedded calcite-cemented sandstones. At the start of the cliff face, these clastic rocks abruptly grade into algal laminated and peloidal limestones (Figs. 9 and 10). As discussed above, Hickson and Lamb sampled a thick ash from within the Bitter Ridge Limestone at this locality (Figs. 9 and 10F) that yielded an age of 14.32 ± 0.10 Ma (single crystal laser fusion of 12 sanidine crystals). This new date suggests that the Bitter Ridge Limestone may be strongly time transgressive from east to west (i.e., younging from east to west). Çakir et al. (1998), while ignoring mapped exposures of the Bitter Ridge Limestone in the Rainbow Gardens area, observed that the Bitter Ridge Limestone (1) outcrops only north of the Las Vegas Valley shear zone and Lake Mead fault system, (2) shows apparent abrupt lateral facies changes from conglomerate into lacustrine limestone facies in the Callville Bay area, and (3) ranges from 13.2 to 13.5 Ma in age. They proposed that an extensive Bitter Ridge Limestone lake system formed north of the Las Vegas Valley shear zone and Lake Mead fault system, with active uplifts to the south of this lake shedding coarsegrained detritus into its shoreline. In their model, the Bitter Ridge Limestone units were inverted post 13.2 Ma due to north-south directed shortening, creating the distinctive contractional structures that we will examine at Stop 3.3. The new age data do not rule out this interpretation, but they do suggest that the Bitter Ridge Limestone lake system was longer-lived and may have had a more complex history than is suggested by Çakir et al.’s (1998) model. If the Bitter Ridge Limestone is older in the east and youngs significantly to the west, this may imply a westward shift in the depocenter due to tectonic processes. The Thumb and Bitter Ridge Limestone strata at this locality contrast strongly with the units examined so far (Fig. 9). In Echo Wash, east of Stop 2.7, there is a thick gypsiferous sequence overlying Thumb Member siliciclastic units that may be correlative with the Bitter Ridge Limestone. Hickson and Lamb have sampled a thick tuff unit in this evaporitic sequence for both geochronological and geochemical analysis in an effort to ascertain whether this hypothesized correlation is valid. Cumulative mi (km) 0.0 0.6

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Directions Continue driving up the wash. Stop 2.4 is at the northern mouth of the canyon between West Longwell Ridge and Bitter Ridge. The upper part of the Bitter Ridge Limestone Member is to the west.

Development of Miocene faults and basins in the Lake Mead region Stop 2.4: Bitter Spring Quadrangle—Teepee Structures in the Upper Bitter Ridge Limestone Purpose To examine teepee structures in Bitter Ridge Limestone algal laminites. Comments On the west side of the road, several spectacular, large scale (1.5 m wide by 1 m tall) teepee structures are exposed near the top of the Bitter Ridge Limestone Member (Fig. 10G). These structures, commonly encountered in algal laminated carbonates, suggest early cementation of the substrate, followed by desiccation and polygonal fracturing. Subsequent groundwater “pumping” due to flooding of the cemented substrate leads to overthrust and upwarped geometries. Note also the thick exposure of the Bitter Ridge Limestone. Comparable structures have been found near the top of the Bitter Ridge Limestone face exposed in Figure 10F, at the top of section 8 (Fig. 9), suggesting a possible correlation, but this relationship is highly speculative given the numerous small structures exposed in the canyon to the south. Cumulative mi (km) 0.0 0.3

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Directions Continue driving up the wash. Turn right onto a side dirt road and continue for a few hundred meters to the large White Basin fault exposure against the base of the steep cliff to the east. Stop 2.5.

Stop 2.5: White Basin, Bitter Spring Quadrangle—White Basin Fault Exposure Purpose To examine a world-class exposure of the White Basin fault and see evidence of one of the latest stages of extension. Comments At this stop, we will see a beautifully exposed fault surface with slickenlines, fault gouge, and a polished surface (Fig. 10H). Deformed Lovell Wash Member layers are juxtaposed against Paleozoic carbonates. Slickenlines on this fault surface suggest that is a left-lateral normal oblique fault (trend and plunge of slickenlines: 250°, 46°) similar to faults measured 0.5–1.0 km to the east. Çakir et al. (1998) and Duebendorfer and Simpson (1994) have suggested that the latest stage of deformation in this area is faulting along northeast-southwest–striking, left-lateral normal oblique faults. Estimates of magnitude of faulting suggest that this last episode produced smaller offsets than the earlier stages and does not represent the majority of Miocene faulting. Cumulative mi (km) 0.0

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Directions Turn around and backtrack to Echo Wash.

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Turn right onto main Echo Wash road. Stop 2.6.

Stop 2.6: White Basin, Bitter Spring Quadrangle—Lovell Wash Member of the Horse Spring Formation Purpose To examine the uppermost member of the Horse Spring Formation, the Lovell Wash Member in the White basin. Comments This is a representative exposure of the sandstone, tuffaceous sandstone, and tuff of the Lovell Wash Member. Cumulative mi (km) 0.0

(0.0)

1.0

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Directions Retrace Echo Wash road back to the intersection with the Bitter Spring Road. A large sign at this intersection indicates mileages to Buffington Pockets and other destinations. Reset odometer and continue down Echo Wash. At this distance from the intersection is Stop 2.7.

Stop 2.7: Echo Wash, Bitter Spring Quadrangle—Thumb Member (Optional) Purpose To examine younger Thumb sandstone units. Comments Deposition in the Thumb Member abruptly shifts from carbonates to tuffaceous sandstones and mudstones east of Stop 2.1. An 40Ar/39Ar date from an ash bed near this locality yields an age of 14.85 ± 0.07 Ma, making these units roughly correlative with the clastic units between the Longwell Ridges and ~500 k.y. older than the Bitter Ridge Limestone. Low-amplitude, small-scale, open folds characterize the Thumb Member in the eastern Echo Wash area, in contrast to the generally unfolded units to the west. Return to Echo Bay Resort; end of Day 2. Day 3: Callville Bay Quadrangle Stops for this day will concentrate on the part of the Frenchman-Overton corridor within the Callville Bay quadrangle mapped by Anderson (2003; Fig. 11). Structurally, the quadrangle is located in an area where strains associated with the Las Vegas Valley shear zone and Lake Mead fault system conjoin and are focused. Rocks of the Horse Spring Formation are extensively exposed in the north part of the quadrangle, whereas the south part is dominated by Miocene volcanic rocks of the Hamblin lobe of the Hamblin-Cleopatra volcano and a group of satellitic volcanics to the west we call the Callville volcanic series. For the purpose of discussing the Miocene sedimentary rocks, the north part is divided into three structural blocks, which, from

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west to east, are the West End, Lovell, and Callville, named for the washes that drain them. A block of Mesozoic rock between the Lovell and Callville blocks is called the Bowl of Fire block (Fig. 11). The Lovell block contains the type area of the Lovell Wash Member of the Horse Spring Formation, possibly 0.5 km thick, resting conformably on magnificently exposed sections of Bitter Ridge Limestone and Thumb Members (~1.2 km thick). Gravel is very sparse in the Thumb rocks of the Lovell block but abundant in the West End and Callville blocks. The minor gravel in the Lovell block contains clasts suggestive of a southern provenance, totally unlike the abundant gravels in the West End and Callville blocks. The four blocks are bounded by north- to northeast-striking faults that are almost wholly younger than the Horse Spring Formation. Anderson (2003) suggests that the Thumb sequences are disparate because they were juxtaposed by these block-bounding faults, all of which have histories of strike slip and strong associated folding. The stops are chosen to elucidate relations between faulting and sedimentation. The last three focus on problems and controversies over the architecture of basins that are younger than the Horse Spring Formation, namely those in which the informal red sandstone unit and younger strata accumulated. We will discuss the tectonic implications of basin architecture in the light of major and minor strike-slip faults and other regional late Tertiary strain features and evaluate the impact on choices of tectonic models. Cumulative mi (km) 0.0 4.8 6.8 10– 12

(0.0) (8.0) (11.3) (16.7– 20.0)

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Directions Leave Echo Bay Resort. Turn left onto North Shore Road. Cross Echo Wash. View Echo Hills to the east and north (right side). These are a pop-up structure at a right bend in the left-lateral Bitter Springs Valley fault system (Campanga and Aydin, 1991). Pass Bitter Springs road, the road into the Longwell Ridge–White basin area. Callville Wash, Stop 3.1. (See additional walking directions below.)

Stop 3.1: Lovell Wash, Callville Bay Quadrangle— Conglomerate Facies of the Lovell Wash and Bitter Ridge Limestone

left-sense slip, opposite to the Las Vegas Valley shear zone. Continue walking up the wash along strike of well-exposed, coarse fault-front debris deposits that dip to the northeast ~50°. Within 100–200 m along strike, observe the bedding-parallel facies transition from the coarse debris through interbedded gravel and sand to thinly laminated mudstone, limey mudstone, algal limestone, and sandstone, all of lacustrine origin. The coarse debris spans the contact between the Bitter Ridge Limestone and Lovell Wash Members of the Horse Spring Formation as mapped by Bohannon (1983a) and Anderson (2003). It consists mostly of upper Paleozoic carbonate rocks and Mesozoic clastic rocks that were derived from a southerly or southeasterly source. Duebendorfer et al. (1998) suggested that this debris was derived from the Frenchman Mountain block (now located ~35 km west of here), which they restored to a position adjacent to these deposits between 13.5 and 12.5 Ma. A debris deposit of similar thickness and much greater volume than the one seen here is exposed over an east-west strike length of almost 10 km ~1.3 km south of here. It is mapped with the Thumb Member (Anderson, 2003) but its age is not well constrained. The debris in that deposit was derived from the north. It might also be a candidate for recording the tectonic transport of a large structural block such as the Frenchman Mountain block. If so, the slightly older time would require restoration to a position west of the deposits at this stop. Improving the age constraints and provenance of coarse fault-front deposits and landslide blocks in the Frenchman-Overton corridor should be a high priority for future basin studies because of the potential for constraining fault juxtapositions. Cumulative mi (km) 0.0 2.8

(0.0) (4.7)

Directions Return to pavement and turn right (west). West End Wash turn off. Turn right onto the dirt road. There will be multiple hairpin turns in the short distance to the wash bottom.

Stop 3.2: West End Wash, Callville Bay Quadrangle— Rainbow Gardens Member on Jurassic Aztec Unconformity

Purpose To look at the abrupt facies transition in upper Horse Spring Formation exposed in Lovell Wash.

Purpose To compare the lithology of a well-exposed tilted section of basal conglomerate and overlying limestone of the Rainbow Gardens Member with the lithology of equivalent beds seen at Stop 1.2. The red sandstone unit and the Muddy Creek Formation will also be viewed.

Comments Walk 700 m northwest to the east lip of Lovell Wash, descend into the wash, and walk up the wash across the trace of a major east-west–striking fault zone. This fault zone has been referred to as the easternmost extent to the Las Vegas Valley shear zone. Here, fault fabrics and kinematic indicators reveal

Comments Here, the Rainbow Gardens rocks rest unconformably on reddish-orange rocks of the Jurassic Aztec Sandstone. This exposure of the basal Rainbow Gardens is important because it is the only one known for 20 km to the west and 7 km to the east. The rocks here form the southwest corner of a pod-form structural

Development of Miocene faults and basins in the Lake Mead region block of south- and southwest-dipping Permian through Jurassic strata we refer to as the Highway block because the North Shore Road passes through it. Equivalent Rainbow Gardens–Aztec contacts to the one seen here are mapped at the northeast corner of the Frenchman Mountain block 20 km west and at the southwest end of the Razorback Ridge block 7 km east. All three localities are located in the structural domain between the large-displacement Las Vegas Valley shear zone on the north and strands of the Lake Mead fault system on the south and southeast. We speculate that the rocks at the three localities were once more proximal to one another than they are now, perhaps even connected to one another. Because the three localities are in a single structural domain, speculation about their former proximity raises questions about the magnitude and style of internal strain on unrecognized structures within any particular structural domain in this region. Cumulative mi (km) 0.0

(0.0)

~0.3

(0.5)

~1.1

(1.8)

Directions Drive up (north) along the West End Wash from this locality. At first, wash cuts reveal gently south-dipping gravel and sand of a syntectonic basin-fill assemblage we refer to as the red sandstone unit. This unit ranges in age from ca. 11.4–8.5 Ma. Throughout the Lake Mead region, rocks of this age, which include a series of andesites, basalts, and interstratified volcaniclastic rocks to the south and southeast of here, record development of syntectonic basins. Controversy exists as to whether those basins record large-magnitude east-west extension, north-south shortening, or both. Along the east wall of the wash, the red sandstone unit is overlain unconformably by gravel of the younger Muddy Creek Formation. The western wash cuts reveal an unconformity between the gravel of the red sandstone unit and 50°–60° south-dipping Triassic rocks. About 100 m north of those exposures, where a tributary wash enters West End Wash from the west, the wash crosses the east-west trace of the Las Vegas Valley shear zone. Gap between the imposing cliffs formed in south-dipping yellowish gray limestone of the Bitter Ridge Limestone and Stop 3.3.

Stop 3.3: West End Wash, Callville Bay Quadrangle—Thumb, Bitter Ridge Limestone, Red Sandstone Unit, and Structure Purpose The main purpose of this stop is to consider the influence of the Las Vegas Valley shear zone on the location, orientation, and deformation of a syntectonic depocenter. The stop is divided into two parts.

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Comments for Part 1 This part of the stop is multipurpose to view (1) the contact between Thumb Member and overlying Bitter Ridge Limestone, (2) the uniform lithology of the Bitter Ridge Limestone Member of the Horse Spring Formation, (3) the sheared upper contact between the Bitter Ridge Limestone and a thin selvage of overlying Lovell Wash Member, and (4) the north-derived provenance of locally derived angular clasts of Bitter Ridge Limestone in steeply south-dipping gravel of the red sandstone unit. The steep north-facing slope to the left of the gap reveals the abrupt upward transition from clastic rocks of the Thumb Member to lacustrine limestone of the Bitter Ridge Limestone Member. This same contact is more distant west of the gap owing to a leftlateral fault that passes through the gap. This contact zone has a uniform appearance over a broad (70 km2) area east and northeast of here, lending confidence to correlation over that area. Walk down-wash and observe the uniform lithology of the Bitter Ridge Limestone. As has been noted previously on this field trip, algal limestones such as these have a considerable age range, introducing much skepticism about correlations from area to area. In the cliffs west of the gap, there are numerous striated fault surfaces in the limestone that support the presence of a northeaststriking fault passing through the gap. The limestone east of the gap is folded, possibly in association with the faulting. Continue walking south along the west wash margin and observe the steeply south-dipping and somewhat sheared-out contact between the Bitter Ridge Limestone and a thin selvage of overlying light gray to whitish tuffaceous sandstone possibly belonging to the Lovell Wash Member of the Horse Spring Formation. For the next 100 m or so to the south, discontinuous exposures of vertical-dipping gravel, some containing angular clasts of Bitter Ridge Limestone, can be seen along the margin of the wash. These beds are mapped as the syntectonic red sandstone unit. They dip more steeply than the underlying Bitter Ridge Limestone, a structural aspect consistent with contraction across a steep reverse fault separating the two units. Thus, although we have passed southward into a syntectonic basin assemblage, the north basin margin has been foreshortened by north-south contraction. Continue walking south and climb onto the west wash wall to observe additional exposures of the steep, south-dipping to overturned red sandstone unit. On return to the parking location, turn right (west) and walk up the tributary drainage cut into red sandstone and mudstone of the Lower Red Member of the Moenkopi Formation. Comments for Part 2 This part of the stop is also multipurpose to view (1) the Las Vegas Valley shear zone, (2) evidence for left-lateral slip on the shear zone here, (3) related fold structures, and (4) south-derived local provenance of clasts in the strongly folded red sandstone unit. The tributary wash follows the trace of the Las Vegas Valley shear zone. The shear zone is exposed in the right (north) wash bank ~0.5 km up the wash where it separates mudstones of the lower red Member of the Moenkopi Formation on the south from

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similarly colored red gravel rich in clasts of the adjacent Moenkopi. Although this gravel differs dramatically from that observed at Part 1 of this stop (due to contrasting provenance), they are both mapped with the informal red sandstone unit. In this part of the wash, note that the Moenkopi beds in the south wash wall are overlain unconformably by gravel. That gravel is also mapped with the informal red sandstone unit, although it is clearly younger than the gravel that is in fault contact with the Moenkopi. Also note that the Moenkopi beds here dip much more gently than those we observed while driving up the wash. In fact, as we walk up the wash, note that the Moenkopi beds pass through a fold hinge and dip gently north, defining a shallow east-west anticline directly south of the trace of the Las Vegas Valley shear zone. Continue up-wash to ~1 km from the wash mouth (to UTM 0702647E; 4008642N) where a splay from the main fault is exposed at the base of the north wash wall. Shear fabric here is consistent with left-sense displacement as it is at the main fault 100 m to the southwest. Beds of the red sandstone unit here dip 80° N-NW and are probably overturned. The core of the fault zone is occupied by structurally exotic (structurally high) sandstone probably belonging to the Permian Esplanade Formation. Walk north up a steep gully to the western lip of the colluvial apron (at UTM 0702614E; 4008662N) where a view to the west reveals the Permian (?) sandstone core of a tight anticline developed along the splay fault in gravely mudstone and sandstone of the Miocene red sandstone unit. The axial trace of this anticline strikes ~N50E and is probably more consistent with right-lateral than left-lateral displacement on the Las Vegas Valley shear zone. The depocenter that received the red sandstone unit north of the Las Vegas Valley shear zone is interpreted to be controlled by fault-parallel folding. The clastic debris shed into it was derived from directly adjacent units to the north and south. Subsequent to deposition, the fold-controlled basin continued to contract in north-south compression as revealed by reverse faulting and steep tilting at Part 1 of this stop and strong folding here. Clearly, the intensity of deformation here is related to the proximity to the Las Vegas Valley shear zone, but controversy exists as to whether north-south shortening strains are restricted to a narrow zone along such faults or are distributed regionally. We will revisit this subject on the last stop. Cumulative mi (km) 0.0

0.8

(0.0)

(1.3)

Directions Return to vehicles. Drive down (south) along the wash. On the east (left) bank, ~0.7 km from the last stop, note a small (m-scale) south-facing monocline developed in red sandy mudstone of the red sandstone unit. These small-scale easterly trending folds are common in Tertiary sedimentary rocks in the area and generally have reverse faults in their core zones. They record northsouth shortening, consistent with the strain recorded in map-scale structures. Intersection with North Shore Road; turn right.

1.5

(2.5)

Turn left onto the dirt road and drive 0.2 mi to the point where the road intersects a power line and park.

Stop 3.4: Pole Line, Callville Bay Quadrangle—Callville Volcanics, Muddy Creek Formation, Basin Architecture Purpose To view basin-fill sediments of the Muddy Creek Formation (ca. 8.5–5.8 Ma) and consider the structural nature of depocenters in which the Muddy Creek in this area accumulated. Comments Across the wash from the power line, basalt flows at the top of the Callville volcanic sequence (time equivalent of the red sandstone unit) dip ~20° north and are overlain by gently north-dipping sand and pebbly sand. The sand could be part of an eolian sand ramp plastered onto the basalt paleoslope. Walk north up-wash to observe (1) a coarsening upward basin-fill sedimentary sequence, (2) clast imbrications indicating south-directed transport, and (3) a reversal in dip direction as a synclinal axis is crossed. These basin-fill sediments are interpreted to be captured in a west-widening, east-west–trending depocenter developed by downwarping in the pathway of south-draining paleodrainages. Alternatively, they are simply folded sediments. This is the most accessible of several such “depocenters” of Muddy Creek age in this part of the Frenchman-Overton corridor. They are similar in orientation to the basin that received the red sandstone unit at the previous stop, but are less deformed and are not in close proximity to major faults. This “depocenter,” one 3 km to the east, and another 5 km to the southeast, lack normal faults along their east margins and thus have little, if any, relationship to east-west extension. To the south of this locality, upper flows of the Callville volcanic sequence are folded on east-west axes into broad open folds. To the southeast, early flows of the sequence are interbedded with coarse volcanogenic debris, cut by reverse faults, and capped at two stratigraphic levels by internal angular unconformities, all providing evidence of strong synvolcanic north-south to northeast-southwest shortening. Thus, to the north and south of here, structures suggest protracted north-south shortening in strata immediately older than the Muddy Creek. Deposition of the Muddy Creek in a north-south contractional setting is not anomalous. Cumulative mi (km) 0.0

(0.0)

1.4

(2.3)

7.2

(12.0)

8.2

(13.7)

Directions Return to vans. Turn left (west) onto Northshore Road. Turn left on marked road to Callville Bay Marina. View Muddy Creek Formation on left side of road. Turn right onto campground access road and keep left (straight) into lower campground and park.

Development of Miocene faults and basins in the Lake Mead region Stop 3.5: Callville Bay Campground, Callville Bay Quadrangle—Fold in Roundstone Gravel—Pliocene? Comments Walk south through the campground onto a graded dirt service road to the first wash bottom to view weakly consolidated roundstone gravel and overlying moderately well sorted sands with sparse suspended pebbles (Fig. 12). These beds record deposition following the establishment of through-flowing latest Tertiary or earliest Quaternary drainages precursor to the Colorado River drainage system. The beds here dip ~40° N. We will walk north up the wash to observe similar, and probably equivalent, beds that dip ~30° S, showing that these strata are folded on approximately east-west axes. Additional folds with somewhat lesser limb dips are mapped to the south in similar strata (Anderson, 2003). Also, Longwell (1936) reported similar structures in river gravels currently beneath the waters of Lake Mead. The style and youthfulness of these folds suggest that contractional strain similar in origin to that which dominates the structural fabric of this part of the Frenchman-Overton corridor is part of the neotectonic regime. Also, rather than being restricted to narrow zones along strike-slip faults such as the Las Vegas Valley shear zone, late Miocene easttrending folds and associated reverse faults and northerly striking strike-slip faults are distributed in a north-south belt at least 15 km long at this longitude (Anderson, 2003). Whether or not these strains are important to tectonic reconstructions is debatable. If they are added to shortening associated with steep-axis rotations in this part of the Frenchman-Overton corridor, the sum of the shortening could exceed 10 km. To date, these strain elements and others such as the possible southward displacement of the Frenchman Mountain block noted at Stop 1.4 have not been factored into tectonic reconstructions. End of field trip. REFERENCES CITED Anderson, R.E., 1971, Thin skin distension in Tertiary rocks of southeastern Nevada: Geological Society of America Bulletin, v. 82, p. 43–58. Anderson, R.E., 1973, Large-magnitude Late Tertiary strike-slip faulting north of Lake Mead, Nevada: U.S. Geological Survey Professional Paper 0794, 18 p. Anderson, R.E., 1977a, Composite stratigraphic section of Tertiary rocks in the Eldorado Mountains, Nevada: U.S. Geological Survey Open-File Report 77-483, 5 p. Anderson, R.E., 1977b, Geologic map of the Boulder City 15-minute quadrangle, Clark County, Nevada: U.S. Geological Survey, scale 1:62,500. Anderson, R.E., 1978a, Chemistry of Tertiary volcanic rocks in the Eldorado Mountains, Clark County, Nevada: U.S. Geological Survey Journal of Research, v. 6, no. 3, p. 409–424. Anderson, R.E., 1978b, Geologic Map of the Black Canyon 15-minute quadrangle, Mohave County, Arizona, and Clark County, Nevada: U.S. Geological Survey Map GQ-1394, scale 1:62,500. Anderson, R.E., 2003, Geologic map of the Callville Bay quadrangle, Clark County, Nevada, and Mohave County, Arizona: Nevada Bureau of Mines and Geology, scale 1:24,000. Anderson, R.E., and Barnhard, T.P., 1993, Aspects of three-dimensional strain at the margin of the extensional orogen, Virgin River depression area, Nevada, Utah, and Arizona: Geological Society of America Bulletin, v. 105, p. 1019– 1052, doi: 10.1130/0016-7606(1993)105<1019:AOTDSA>2.3.CO;2. Anderson, R.E., Barnhard, T.P., and Snee, L.W., 1994, Roles of plutonism, midcrustal flow, tectonic rafting, and horizontal collapse in shaping the

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A

B

Figure 12. Photographs of key outcrop features from Day 3 field stops. (A) Tilted, moderately well sorted quartzose sandstone directly south of Callville Bay campground. These beds are interbedded with roundstone gravel and are mapped as Quaternary-Tertiary alluvium by Anderson (2003). They are tilted in the north limb of a syncline, the axis of which passes through the campground. Although their age is not known, they are unlikely to be older than 4 Ma. (B) View west along north shore of Lovers Cove southwest of the Callville Bay campground showing 20° north-tilted interbedded sands and roundstone gravels mapped as Quaternary-Tertiary mainstream alluvium. Tilting is in the north limb of an anticline. The roundstone gravels include quartzite and chert and represent deposition by through-flowing streams precursor to the Colorado River.

Miocene strain field of the Lake Mead area, Nevada and Arizona: Tectonics, v. 13, no. 6, p. 1381–1410, doi: 10.1029/94TC01320. Anderson, R.E., Longwell, C.R., Armstrong, R.L., and Marvin, R.F., 1972, Significance of K-Ar ages of Tertiary rocks from the Lake Mead region, NevadaArizona: Geological Society of America Bulletin, v. 83, p. 273–287. Axen, G.J., Taylor, W.J., and Bartley, J.M., 1993, Space-time patterns and tectonic controls of Tertiary extension and magmatism in the Great Basin of the Western United States: Geological Society of America Bulletin, v. 105, p. 56–76, doi: 10.1130/0016-7606(1993)105<0056:STPATC>2.3.CO;2. Axen, G.J., Wernicke, B.P., Skelly, M.F., and Taylor, W.J., 1990, Mesozoic and Cenozoic tectonics of the Sevier thrust belt in the Virgin River valley area, southern Nevada, in Anderson, J.L., ed., Cordilleran magmatism: Geological Society of America Memoir 176, p. 123–153. Beard, L.S., 1996, Paleogeography of the Horse Spring Formation in relation to the Lake Mead fault system, Virgin Mountains, Nevada and Arizona, in Beratan, K.K., ed., Reconstructing the history of Basin and Range extension using sedimentology and stratigraphy: Geological Society of America Special Paper 303, p. 27–60.

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Beard, L.S., Anderson, R.E., Block, D.L., Bohannon, R.G., Brady, R.J., Castor, S.B., Duebendorfer, E.M., Faulds, J.E., Howard, K.A., Kuntz, M.A., Rowlands, S.M., Wallace, M.A., and Williams, V.S., 2006, Geologic map of the Lake Mead 30′ × 60′ Quadrangle, Nevada: U.S. Geological Survey Scientific Investigations Map Series, scale 1:100,000 (in press). Bohannon, R.G., 1979, Strike-slip faults of the Lake Mead region of southern Nevada, in Armentrout, J.M., Cole, M.R., and Terbest, H., Jr., eds., Pacific Coast Paleogeography Symposium no. 3: Pacific Section, Society of Economic Paleontologists and Mineralogists (SEPM), p. 129–139. Bohannon, R.G., 1983a, Geologic map, tectonic map and structure sections of the Muddy and northern Black Mountains, Clark County, Nevada: U.S. Geological Survey, scale 1:62,500. Bohannon, R.G., 1983b, Mesozoic and Cenozoic tectonic development of the Muddy, North Muddy, and northern Black Mountains, Clark County, Nevada, in Miller, D.M., Todd, V.R., and Howard, K.A., Tectonic and stratigraphic studies in the eastern Great Basin: Geological Society of America Memoir 157, p. 125–148. Bohannon, R.G., 1984, Nonmarine sedimentary rocks of Tertiary age in the Lake Mead region, southeastern Nevada and northwestern Arizona: U.S. Geological Survey Report P-1259, 72 p. Çakir, M., Aydin, A., and Campagna, D., 1998, Deformation pattern around the conjoining strike-slip fault systems in the Basin and Range, southeast Nevada: the role of strike-slip faulting in basin formation and inversion: Tectonics, v. 17, no. 3, p. 344–359, doi: 10.1029/98TC00562. Campagna, D.J., and Aydin, A., 1991, Tertiary uplift and shortening in the Basin and Range; the Echo Hills, southeastern Nevada: Geology, v. 19, p. 485– 488, doi: 10.1130/0091-7613(1991)019<0485:TUASIT>2.3.CO;2. Campagna, D.J., and Aydin, A., 1994, Basin genesis associated with strike-slip faulting in the Basin and Range, southeastern Nevada: Tectonics, v. 13, no. 2, p. 327–341, doi: 10.1029/93TC02723. Castor, S.B., Faulds, J.E., Rowland, S.M., and dePolo, C.M., 2000, Geologic map of the Frenchman Mountain Quadrangle, Clark County, Nevada: Nevada Bureau of Mines and Geology, scale 1:24,000. Donatelle, A.R., Goeden, J., Hannon, M., Hickson, T., Holter, S., Johnson, T., Lamb, M., Lindberg, J., 2004, Implications of stratigraphic and structural data from the Bitter Spring region, Southern Nevada: Eos (Transactions, American Geophysical Union), v. 85, no. 17, Joint Assembly Supplement, Abstract T31B-09. Donatelle, A.R., Hickson, T., and Lamb, M., 2005, Reevaluation of Miocene stratigraphy in light of new stratigraphic, petrographic and 39Ar/40Ar data, north of Lake Mead, Nevada: Geological Society of America Abstracts with Programs, v. 37, no. 5, p. 54. Duebendorfer, E.M., 2003, Geologic map of the Government Wash Quadrangle, Clark County, Nevada: Nevada Bureau of Mines and Geology, scale 1:24,000. Duebendorfer, E.M., and Black, R.A., 1992, Kinematic role of transverse structures in continental extension; an example from the Las Vegas Valley shear zone, Nevada: Geology, v. 20, p. 1107–1110, doi: 10.1130/00917613(1992)020<1107:KROTSI>2.3.CO;2. Duebendorfer, E.M., and Sharp, W.D., 1998, Variation in displacement along strike of the South Virgin-White Hills detachment fault; perspective from the northern White Hills, northwestern Arizona: Geological Society of America Bulletin, v. 110, p. 1574–1589, doi: 10.1130/00167606(1998)110<1574:VIDASO>2.3.CO;2. Duebendorfer, E.M., and Simpson, D.A., 1994, Kinematics and timing of Tertiary extension in the western Lake Mead region, Nevada: Geological Society of America Bulletin, v. 106, p. 1057–1073, doi: 10.1130/00167606(1994)106<1057:KATOTE>2.3.CO;2. Duebendorfer, E.M., and Wallin, E.T., 1991, Basin development and syntectonic sedimentation associated with kinematically coupled strike-slip and detachment faulting, southern Nevada: Geology, v. 19, p. 87–90, doi: 10.1130/0091-7613(1991)019<0087:BDASSA>2.3.CO;2. Duebendorfer, E.M., Beard, L.S., and Smith, E.I., 1998, Restoration of Tertiary deformation in the Lake Mead region, southern Nevada; the role of strikeslip transfer faults, in Faulds, J.E., and Steward, J.H., eds., Accommodation zones and transfer zones; the regional segmentation of the Basin and Range Province: Geological Society of America Special Paper 323, p. 127–148. Eaton, G.P., Wahl, R.R., Prostka, H.J., Mabey, D.R., and Kleinkopf, M.D., 1978, Regional gravity and tectonic patterns; their relation to late Cenozoic epeirogeny and lateral spreading in the western Cordillera, in Smith, R.B., and Eaton, G.P., eds., Cenozoic tectonics and regional geophysics of the Western Cordillera: Geological Society of America Memoir 152, p. 51–91. Feuerbach, D.L., Smith, E.I., Shafiqullah, M., and Damon, P.E., 1991, New KAr dates for late to early Pliocene mafic volcanic rocks in the Lake Mead area, Nevada and Arizona: Isochron West, no. 57, p. 17–20.

Fryxell, J.E., and Duebendorfer, E.M., 2005, Origin and trajectory of the Frenchman block, an extensional allochthon in the basin and range province, southern Nevada: Journal of Geology, v. 113, p. 355–372, doi: 10.1086/428810. Harlan, S.S., Duebendorfer, E.M., and Deibert, J.E., 1998, New 40Ar/39 Ar isotopic dates from Miocene volcanic rocks in the Lake Mead area and southern Las Vegas Range, Nevada: Canadian Journal of Earth Sciences (Revue Canadienne des Sciences de la Terre), v. 35, no. 5, p. 495–503. Langenheim, V.E., Grow, J.A., Jachens, R.C., Dixon, G.L., Miller, J.J., Lundstrom, S.C., and Page, W.R., 2001, Basin configuration beneath Las Vegas Valley, Nevada; implications for seismic hazard evaluation, in Luke, B.A., Jacobson, E.A., and Werle, J.L., eds., Proceedings of the Symposium on Engineering Geology and Geotechnical Engineering, v. 36, p. 755–764. Liggett, M.A., and Childs, J.F., 1977, An application of satellite imagery to mineral exploration, U.S. Geological Survey Professional Paper P-1015, p. 253–270. Longwell, C.R., 1928, Geology of the Muddy Mountains, Nevada, with a section through the Virgin Range to the Grand Wash Cliffs, Arizona, U.S. Geological Survey Bulletin, Report 0798, 152 p. Longwell, C.R., 1936, Geology of the Boulder reservoir floor, Arizona-Nevada: Geological Society of America Bulletin, v. 47, p. 1393–1476. Longwell, C.R., 1974, Measure and date of movement on Las Vegas Valley shear zone, Clark County, Nevada: Geological Society of America Bulletin, v. 85, p. 985–989, doi: 10.1130/0016-7606(1974)85<985:MADOMO>2.0.CO;2. Longwell, C.R., Pampeyan, E.H., Bowyer, B., and Roberts, R.J., 1965, Geology and mineral deposits of Clark County, Nevada: Bulletin, Nevada Bureau of Mines, 218 p. Martin, L., 2005, Miocene stratigraphy and sedimentology in Longwell Ridges area, and implications for extension, Lake Mead region, Nevada [M.S. thesis]: Flagstaff, Northern Arizona University, 155 p. Pederson, J.L., Pazzaglia, F.J., Smith, G.A., and Mou, Y., 2000, Neogene through Quaternary hillslope records, basin sedimentation, and landscape evolution of southeastern Nevada, in Lageson, D.R., ed., and Peters, S.G., and Lahren, M.M, coeditors, Great Basin and Sierra Nevada: Geological Society of America Field Guide 2, p. 117–134. Price, L.M., 1997, The geometry and evolution of a major segment of the Grand Wash fault zone and associated growth-fault basin, southern White Hills, northwestern Arizona: [M.S. thesis]: Iowa City, University of Iowa, 148 p. Ron, H., Aydin, A., and Nur, A., 1986, Strike-slip faulting and block rotation in the Lake Mead fault system: Geology, v. 14, p. 1020–1023, doi: 10.1130/ 0091-7613(1986)14<1020:SFABRI>2.0.CO;2. Rowland, S.M., Parolini, J.R., Eschner, E., McAllister, A.J., and Rice, J.A., 1990, Sedimentologic and stratigraphic constraints on the Neogene translation and rotation of the Frenchman Mountain structural block, Clark County, Nevada, in Anderson, J.L., ed., Cordilleran magmatism: Geological Society of America Memoir 176, p. 99–122. Snow, J.K., and Wernicke, B., 2000, Cenozoic tectonism in the central Basin and Range; magnitude, rate, and distribution of upper crustal strain: American Journal of Science, v. 300, no. 9, p. 659–719. Sonder, L.J., Jones, C.H., Salyards, S.L., and Murphy, K.M., 1994, Vertical axis rotations in the Las Vegas Valley shear zone, southern Nevada; paleomagnetic constraints on kinematics and dynamics of block rotations: Tectonics, v. 13, no. 4, p. 769–788, doi: 10.1029/94TC00352. Spencer, J.E., and Reynolds, S.J., 1989, Middle Tertiary tectonics of Arizona and adjacent areas, in Jenney, J.P., and Reynolds, S.J., Geologic evolution of Arizona: Arizona Geological Society Digest 17, v. 539–574. Stewart, J.H., 1998, Regional characteristics, tilt domains, and extensional history of the later Cenozoic Basin and Range Province, western North America, in Faulds, J.E., and Stewart, J.H., eds., Accommodation zones and transfer zones; the regional segmentation of the Basin and Range Province: Geological Society of America Special Paper 323, p. 47–74. Thompson, K.G., 1985, Stratigraphy and petrology of the Hamblin-Cleopatra Volcano, Clark County, Nevada [Master’s thesis]: Austin, University of Texas, 306 p. Wawrzyniec, T.F., Geissman, J.W., Anderson, R.E., Harlan, S.S., and Faulds, J.E., 2001, Paleomagnetic data bearing on style of Miocene deformation in the Lake Mead area, southern Nevada: Journal of Structural Geology, v. 23, no. 8, p. 1255–1279, doi: 10.1016/S0191-8141(00)00191-7. Wernicke, B.P., Axen, G.J., and Snow, J.K., 1988, Basin and Range extensional tectonics at the latitude of Las Vegas, Nevada: Geological Society of America Bulletin, v. 100, p. 1738–1757, doi: 10.1130/00167606(1988)100<1738:BARETA>2.3.CO;2.

Printed in the USA

Geological Society of America Field Guide 6 2005

Don R. Currey Memorial Field Trip to the shores of Pleistocene Lake Bonneville Holly S. Godsey Department of Geology and Geophysics, University of Utah, Salt Lake City, Utah 84112, USA Genevieve Atwood Earth Science Education, 30 North U Street, Salt Lake City, Utah 84103, USA Elliott Lips Department of Geography, University of Utah, Salt Lake City, Utah 84112, USA David M. Miller United States Geological Survey, 345 Middlefield Road, Menlo Park, California 94025, USA Mark Milligan Utah Geological Survey, P.O. Box 146100, Salt Lake City, Utah 84114, USA Charles G. Oviatt Department of Geology, Kansas State University, Manhattan, Kansas 66505, USA

ABSTRACT Donald R. Currey spent over two decades researching and exploring relics of ancient Lake Bonneville in the eastern Great Basin. Shoreline and deepwater deposits of Lake Bonneville document coastal processes, lake chemistry, and environmental change during the late Pleistocene and Holocene. This field guide summarizes findings at many of the classic localities researched by Currey and colleagues that contributed to the current understanding of this impressive pluvial lake and its interglacial successor, Great Salt Lake. Subjects include coastal processes at Antelope Island and the Stockton Bar; lake history, chemistry and environmental change at Stansbury Island, the Public Shooting Grounds and Hansel Valley; deltaic depositional processes at Big Cottonwood Canyon, American Fork Canyon and Brigham City; and the relative chronology of glacial and lacustrine deposition at Little Cottonwood Canyon and Bells Canyon. Keywords: Lake Bonneville, delta, climate change, marl, Provo shoreline, Pleistocene, Holocene.

Godsey, H.S., Atwood, G., Lips, E., Miller, D.M., Milligan, M., and Oviatt, C.G., 2005, Don R. Currey Memorial Field Trip to the shores of Pleistocene Lake Bonneville, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 419–448, doi: 10.1130/ 2005.fld006(19). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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TRIBUTE TO DONALD R. CURREY (24 JAN. 1934–6 JUN. 2004) Modified from G. Atwood, 2004; Reprinted with Permission Donald Rusk Currey was born and raised in California, USA, vacationing in the Sierra Nevada and along the paleochannel of the Owens River. Landscapes, the processes that shape them, and their history of climate change, intrigued him early on and challenged him to the end. Currey attended Stanford University as an undergraduate, worked summers surveying topography and in the Alaskan branch of the United States Geological Survey, transferred to the University of Wyoming for reasons of health, and earned his Bachelor of Science (1957) and Master’s (1959) degrees both in geology. His Ph.D. (1969) from the University of Kansas in physical geography reported his research on post–Ice Age climate change in the mountains of southwestern United States. Currey came to the University of Utah in 1970 and joined the faculty of the department of geography. He rose from associate professor to chair of the department and full professor. He established the University of Utah Limneotectonics Laboratory and was instrumental in founding the DIGIT Lab (Digitally Integrated Geographic Information Technologies Laboratory). He chaired the committees of over 30 Master’s students and 15 doctoral students, and advised three post-doctoral students. He mentored and served as a member of numerous graduate student committees within and outside the department of geography. Currey saw lakes as historians of climate change. His research ranged from saline lakes of the central Andes to glacial lakes of the Canadian Shield. He even co-authored work on extraterrestrial evidence of geomorphic processes. Primarily, though, he focused his research on Lake Bonneville sediments and landforms as the key to unraveling the detailed history of climate change of western continental North America. He taught popular introductory courses in physical geography, undergraduate geoexcursions, and graduate courses in geomorphology of lakes, paleolakes, lake basins, and coasts. Awards included the G.K. Gilbert Award (1992) of the Association of American Geographers, the Geographer of the Year Award (1993) of the Utah Geographical Society, and the Superior Research Award (1994) of the University of Utah College of Social and Behavioral Sciences. His professional activities included consulting for engineering firms, industry, and governmental agencies. He generously gave time and expertise to individuals and organizations including Friends of Great Salt Lake and the Utah Geological Survey. In 2002, Friends of Great Salt Lake awarded Currey the first-ever annual Friend of Great Salt Lake Award. Currey was amazing in the field. He was a skilled observer with an outstanding ability to spot and identify natural features. He thought in four dimensions always, everywhere (latitude, longitude, elevation, and time). He enjoyed working with students, and his students loved him and loved working with him. The classes he taught that involved major field trips to Great Salt Lake and to several regions of the Great Basin were legendary. He val-

ued field work and field experience and made the Great Basin and western United States his students’ outdoor laboratory. A cause that absorbed much of his recent attention was to protect critical landforms, primarily Lake Bonneville features, that preserve important earth systems information and are threatened by destruction from mining or urban sprawl. He joined Marjorie Chan of the University of Utah department of geology and geophysics in a National Science Foundation–funded project to identify these features and develop strategies to protect them. Currey, who was never reluctant to introduce new terms into the scientific literature, titled these features “geoantiquities.” Although Currey studied many lakes, Great Salt Lake and its predecessor, Lake Bonneville, received his most tenacious attention. G.K. Gilbert’s work in the nineteenth century defined major characteristics of Great Salt Lake and Lake Bonneville and related them to climate change. Later researchers had added to the story. Currey and the graduate students working with him made tremendous contributions to the understanding of Lake Bonneville and Great Salt Lake. While many of his publications are technical and appear in professional journals, he also published articles and maps intended for the general public and for policymakers. In his final months, Currey had obtained new dates on numerous Lake Bonneville features that he planned to use to develop a revised history of Lake Bonneville. Don Currey will be remembered primarily for (1) his contributions to the understanding of the history of Pleistocene lakes and the use of these histories in the study of how Earth’s climate has varied under natural conditions, and (2) the students he inspired and trained, who have made, and will continue to make, important contributions in numerous areas. He will be greatly missed by past and present students, colleagues, family, and friends (Fig. 1). INTRODUCTION TO LAKE BONNEVILLE AND THE BONNEVILLE BASIN Lake Bonneville was a large late Pleistocene pluvial lake that occupied the eastern Great Basin from ca. 28–10 ka (Fig. 2) (Gilbert, 1890; Benson et al., 1990; Currey, 1990). This region is part of the Basin and Range physiographic province that developed in response to east-west crustal extension beginning ca. 20 Ma. Normal faulting and subsidence caused the formation of several basins separated by uplifted fault blocks. The most significant subsidence occurred at the eastern margin of the Basin and Range province, and it is here that the Bonneville Basin formed. Extension is continuing to modify the landscape in the Basin and Range province today. The Bonneville Basin has been a region of internal drainage for the past 15 m.y., and lakes of varying sizes have occupied the basin throughout most of this time (Currey et al., 1984). Researchers have produced evidence that at least four deep lake cycles have occurred in the Bonneville Basin over the past 500,000 yr; the Bonneville lake cycle being the youngest of these (Gilbert, 1890; Hunt et al., 1953; Morrison, 1965; Scott et al.,

Don R. Currey memorial field trip

Figure 1. Don Currey investigating Lake Bonneville deposits exposed by the 2002 Utah State Capitol renovation. Photo courtesy of Shizuo Nishizawa.

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FIELD TRIP DAY 1—ANTELOPE ISLAND, STOCKTON BAR, TOOELE ARMY DEPOT, STANSBURY GULCH Day 1 begins with an overview of the Bonneville lake cycle at Antelope Island State Park (Fig. 4), including a discussion of historic levels of Great Salt Lake and flooding associated with the 1980s wet cycle. The trip proceeds to the Stockton Bar, an enormous barrier bar and spit complex, to examine the geomorphic expression of the late transgressive and highstand phases of the lake. The trip continues to the Tooele Army Depot (TEAD) to look at highly detailed and well-preserved Provo shoreline

112 o

NEVADA

1983; Currey and Oviatt, 1985; McCoy, 1987; Oviatt et al., 1987; Machette and Scott, 1988; Balch et al., 2005). Lake Bonneville began to form ca. 28 ka when colder and/ or wetter climate conditions during the last ice age caused the lake to rise (Fig. 3). This transgression continued until ca. 23 ka when changing climate conditions caused the lake to undergo a slight regression. For nearly 3000 yr, the level of the lake oscillated near 1370 m above sea level (masl; 4500 ft above sea level [fasl]), forming the Stansbury shoreline, and then began to rise again. Approximately 16 ka, the lake reached the level of the topographic divide near Zenda, Idaho (Fig. 2), and began to overflow. This marks the beginning of the open-basin phase of the lake and the formation of the Bonneville shoreline (Fig. 3) (Burr and Currey, 1988). Hydro-isostatic subsidence of the basin floor caused the lake to transgress to its maximum level ca. 15 ka. At this time, Lake Bonneville was over 300 m deep and covered an area of ~51,000 km2 (Oviatt and Miller, 1997). Rivers emanating from the high mountains to the east provided large amounts of clastic material that was deposited in deltas along the mountain fronts (Lemons et al., 1996). Spits, barriers, bars, and beaches were supplied by alluvium and weathered bedrock in regions with little riverine input. Muds and marls were deposited in more distal, offshore locations (Oviatt and Miller, 1997). Catastrophic failure of the alluvial fan deposits at the threshold ca. 14.5 ka caused a massive flooding of lake waters into the Snake River drainage. This event, called the Bonneville flood, caused a drop in lake level of >100 m and is believed to have occurred in less than one year’s time (Malde, 1968; O’Connor, 1993). Headward erosion by flood waters shifted the drainage divide to the southeast ~3.2 km near Red Rock Pass, Idaho. The lake restabilized at the new bedrock threshold and the Provo shoreline began to form (Fig. 2) (Burr and Currey, 1988). The lake oscillated at or near the level of the Red Rock Pass outlet until a change in climate conditions permanently drove the lake below the threshold. By ca. 11.5 ka, the lake had reached levels comparable to modern Great Salt Lake (Godsey et al., 2005). The lake rose again briefly between ca. 11 ka and 10 ka to form the Gilbert shoreline (Oviatt et al., 2005) and then again declined. This marks the end of the Bonneville lake cycle and the beginning of its successor, Great Salt Lake (Currey et al., 1984).

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Figure 2. Map showing the extent of Lake Bonneville ca. 15,000 14C yr B.P. Locations of the Zenda and Red Rock Pass thresholds indicated with arrows. Dark gray region indicates the extent of Great Salt Lake (after Schofield et al., 2004).

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Figure 3. Chronology of Lake Bonneville showing major shoreline forming events. Ages are in radiocarbon yr B.P.; elevations are adjusted for isostatic rebound of the basin floor (after Oviatt, 1997).

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deposits. Day 1 ends with a stop at Stansbury Gulch to examine the stratigraphy of early transgressive deposits related to the Stansbury shoreline. Directions to Stop 1.1 From Salt Lake City, travel north ~40 km (25 mi) on I-15 and take Exit 335, Antelope Island–Great Salt Lake. Proceed west along Antelope Drive (S.R. 127) to the entrance of Antelope Island State Park. The Gilbert shoreline is crossed at the intersection of Bluff Road and Antelope Drive. At the entrance gate to the state park, stop and reset mileage to zero. Cumulative mi (km) 0.0

(0.0)

6.8

(10.9)

9.2

(14.8)

Directions Drive west from the entrance across the causeway that separates Farmington Bay to the south from Ogden Bay to the north. Once on the island, bear left (south) on the main road at the intersection, follow signs to Buffalo Point and park in the lot by the concessions building. Stop 1.1: Buffalo Point overlook (12, 394336E, 4543207N; NAD83).

Antelope Island Antelope Island is located between Gilbert and Farmington Bays in Great Salt Lake (Fig. 4). The island is a range near the

eastern edge of the Basin and Range physiographic province. Two contrasting geologic units dominate the island’s bedrock: the generally light-colored Cambrian Tintic Quartzite and the darker metamorphic rocks of the Early Proterozoic Farmington Canyon Complex. This contrast of source materials provides opportunities to study sediment transport directions. Shoreline features on Antelope Island are well developed and representative of those of Great Salt Lake and, to a lesser extent, those of Lake Bonneville. The island is spatially isolated from mainland riverine processes; no perennial streams drain the island, although ephemeral drainages and debris flows transport upland material to the modern shorezone. Furthermore, Antelope Island’s status as a state park provides protection from development; therefore, its shorelines are relatively undisturbed and, with park permission, relatively accessible. Don Currey led multiple formal and informal field excursions to Antelope Island and used the well-preserved shorelines of Bonneville, Gilbert, and Great Salt Lake sequences to demonstrate that shorelines are precious records of climate change. Stop 1.1: Buffalo Point Overlook This stop provides a review of the general history of Lake Bonneville and the evolution of the Bonneville Basin. Several Lake Bonneville landforms are visible along the Wasatch Front from Antelope Island. These include the impressive delta of the Weber River (Hill Air Force Base) and the Bonneville level limit of urban development. East-west contrasts of Great Salt Lake can also be observed from Buffalo Point. The eastern por-

Don R. Currey memorial field trip

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Figure 4. Map of the northern Bonneville Basin showing the locations of field trip stops.

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tion of the lake receives virtually all of the freshwater surface drainage and is separated into three bays: Farmington, Ogden, and Bear River Bays (Fig. 4). The western, saltier, portion of Great Salt Lake is divided by a solid-fill railroad causeway into Gunnison Bay to the north and Gilbert Bay to the south (Fig. 4). On a clear day, the railroad causeway can be seen by looking north, just west of the Promontory Mountains.

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Buffalo Bay provides a vantage point to examine the diverse coastal environments of Antelope Island. The island has a variety of shore features because it is impacted by waves driven by wind from all directions (Currey, 1980). Headland and bay features facing west and northwest can be contrasted with those facing east. West- and northwest-facing shores of Antelope Island experience strong storm winds and high-energy wave conditions

H.S. Godsey et al.

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due to longer fetch than those of east-facing shores. The island’s western shore is dominated by erosional headlands, whereas the island’s eastern shore is characterized by accretional landforms that build out into the lake. From Buffalo Point, the four major Lake Bonneville shorelines are visible across White Rock Bay to the south. Here, the Bonneville level is at ~1600 m (~5250 ft), the Provo level is at ~1485 m (~4870 ft), the Stansbury level is at ~1360 m (~4460 ft), and the Gilbert level is at ~1300 m (~4265 ft) (Doelling et al., 1990) (Fig. 5). Holocene shoreline features range from prehistoric shorelines that may be as high as the Gilbert level to offshore, presently submerged, shore features including mudflats interpreted by Currey (1980) as desiccation polygons associated with Holocene lake lowstands. Historic lake fluctuations include two highstands at ~1284 m (4212 ft) reached during the 1860s–1870s and again in 1986–1987 and a lowstand of 1277.5 m (4191 ft) reached in 1963 (Fig. 6). Shoreline features of the 1986–1987 highstand are visible from Buffalo Point and include anthropogenic debris such as timber along the north shore

of White Rock Bay, undercut beach faces partially obscured by windblown sediments of the back of the bay, and vegetation changes along the bay’s south shore. Directions to Stop 1.2 Retrace route down the hill. Cumulative mi (km) 10.4

16.7

10.8

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18.5

Directions Turn left (north) at intersection to Bridger Bay. Veer left off of main road at intersection to Bridger Bay campground. Proceed to end of road and parking lot at west end of campground. Stop 1.2: Bridger Bay Campground (12, 393879E, 4544010N; NAD83).

Stop 1.2: Antelope Island, Bridger Bay, 1986–1987 Highstand Shoreline Features of Great Salt Lake

Bonneville Stansbury Provo Gilbert

Figure 5. Pleistocene and Holocene shorelines, Antelope Island, Great Salt Lake. View is to the south across White Rock Bay.

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The rise of Great Salt Lake to its historic highstand in 1986 and 1987 left shorelines that can be identified by vegetation changes and modern debris that was incorporated into the beach deposits (Fig. 7). Among the items located in the shoreline deposits: a telephone pole, plastic, a bowling ball, railroad ties, tires, and twigs of locally derived organic matter. The 1986–1987 inundation shoreline located farthest inland has been surveyed relative to the gauged still-water lake elevation and studied in terms of coastal processes of shallow closed-basin lakes (Atwood, 2002). The extent to which shoreline expressions deviate from horizontal due to superelevation caused by wave runup and lake set-up diminishes their utility as accurate horizontal datums. However, the extent that patterns of shoreline superelevation result from coastal geomorphic processes increases their potential to be evidence of wave energy and lake surface conditions at the time of shoreline creation. Bridger Bay faces northwest into strong storm winds of Great Salt Lake. Surveyed elevations of 1986–1987 shoreline expressions along Bridger Bay range from within 0.3 m (1 ft) of the 1986–1987 still-water elevation along the shore east of the campground to elevations >2 m (7 ft) above still-water lake elevation at the back of the bay, downslope of the Bridger Bay picnic pavilions.

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1995

Figure 6. Fluctuation of water-surface altitude in Gilbert Bay, Great Salt Lake, 1847 to present (data courtesy of the U.S. Geological Survey).

Return to main road and turn left (north). Road continues past the beach picnic grounds and the visitors center before intersecting the causeway. Return to I-15 from Antelope Island and head south, via Exit 316, I-215 south. Continue south on I-215, to Exit 22A, I-80 west. Continue west on I-80. Notice the welldefined shorelines on the Oquirrh Mountain salient as we pass into Tooele County. Take Exit 99, Tooele. Reset mileage to zero

Don R. Currey memorial field trip

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at intersection with Saddleback Road and the Conoco Flying J gas station. Cumulative mi (km) 0.0

(0.0)

12.2

(19.6)

13.4

(21.6)

16.1

(25.9)

16.4

(26.4)

16.5

(26.6)

Directions Continue south on S.R. 36 through the town of Tooele. The horizontal surface SE of the highway is an erosion platform that formed during the transgressive and highstand phases of Lake Bonneville. Unconsolidated sediments were eroded and transported by longshore processes and deposited ~6 km to the south, forming the enormous baymouth barrier and spits that make up the Stockton Bar. The white-gray rocks visible on the flank of Two O’Clock Hill are a Tertiary-aged igneous intrusion. Pull off onto dirt road NW of highway, just before the double-pole telephone lines. Veer left onto faint double track road. Proceed slowly uphill to the radio relay tower. Turn around and park near tower. Stop 1.3: Radio Relay Tower (12, 385131E, 4480459N; NAD83).

The Stockton Bar The Stockton Bar is an enormous barrier bar and spit complex that contains some of the most detailed and well preserved records of paleolake history found in the Bonneville Basin. G.K. Gilbert first documented this area in his 1890 monograph, Lake Bonneville (Fig. 8). Don Currey and his colleagues later surveyed this region in detail and provided a comprehensive description of

1986-1987 Shoreline

1980s Lagoon

Older Shore Materials

Figure 7. Depositional expression of 1986–1987 Great Salt Lake shoreline at Bridger Bay, Antelope Island. Stratigraphic position and the incorporation of trash in the shoreline sediments date it to 1986–1987.

the geomorphology of the Stockton Bar as it relates to major lake events (e.g., Burr and Currey, 1988). Currey continued to study the Stockton Bar throughout his career and returned to the site often to take samples and make new observations. Stop 1.3: Radio Relay Tower B5 Shoreline and Stockton Spit Changing climate conditions caused Lake Bonneville to oscillate below threshold control several times during the formation of the Bonneville shoreline complex. Hydro-isostatic subsidence continued during most of the subthreshold intervals, resulting in differing elevations of shorelines deposited at the same threshold. Currey recognized the need to distinguish between the various minor shorelines that composed the Bonneville shoreline complex and developed a nomenclature for these features. Major

Figure 8. View to the southeast of The Great Bar at Stockton, Utah, as drawn by G.K. Gilbert (Gilbert, 1890, plate IX); retouched by Holmes (Hunt, 1982, p.170).

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Bonneville shorelines, formed while the lake was at the Zenda threshold, are given the designations B0–B8 (Currey and Burr, 1988). Subthreshold stages (climate controlled) are designated Ba, Bb, and Bc (Currey and Burr, 1988). Deposition of the Stockton Bar began ca. 15.5 ka when transgressing lake waters deposited a series of spits, barriers, and beach ridges in the valley between the Stansbury and Oquirrh Mountains (Fig. 9). Local hydro-isostatic subsidence of the basin floor increased as Lake Bonneville transgressed into Rush Valley and a prograding and aggrading pattern of spit deposition began (Burr and Currey, 1988). Surface drainage from Rush Valley was

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Figure 9. Aerial photograph of the Stockton Bar area ca. 1966, showing transgressive-age shorelines (T), the Bonneville shoreline complex (B1, B3, B5, B6, B8), the Provo shoreline complex (P0, P1, P3, P7, P9), and Rush Valley Provo- and Gilbert-age shorelines (Rp, Rg). Arrows indicate direction of sediment transport during spit formation. Field trip stops at this locality are indicated by numbered stop signs (modified Figure 4 of Burr and Currey, 1988).

eventually blocked as deposition of the barrier bar continued, isolating the lake waters to the south from the main body of Lake Bonneville (Burr and Currey, 1988; Gilbert, 1890; Gilluly, 1929). The lake continued to transgress until it breached the level of the Zenda threshold located far to the north (Fig. 2). This marks the beginning of the open-basin phase of the lake and deposition of the Bonneville shoreline complex (Fig. 3) (Burr and Currey, 1988). Four major stages of threshold control are evident in the Stockton Bar area. Each stage is thought to have been interrupted by a subthreshold event as evidenced by shorelines at other localities (Gilbert, 1890; Burr and Currey, 1988). B1, the first threshold controlled shoreline, is an enormous cross-valley baymouth barrier supplied by sediments primarily from the northeast (Fig. 9). Despite the lowered water plane, hydro-isostatic subsidence continued in this area during a subthreshold interval. Following a return to threshold control, a large 1.5-kmlong spit was deposited southward toward the present-day town of Stockton. This spit, called the Stockton spit, sits at an elevation of 1589 masl (5212 fasl) and represents the B3 shoreline in this region. Following another subthreshold interval and continued subsidence, the lake transgressed again and deposited the B5 spit located at ~1594 masl (5231 fasl) (Fig. 9). This marks the highest elevation of Lake Bonneville attained in the Stockton Bar region (Burr and Currey, 1988). Following the highstand of Lake Bonneville, a noncatastrophic incision of the Zenda threshold caused the lake to drop ~12 m (40 ft). The lake lingered just long enough at this level to winnow the fine sand and gravel matrix from between the cobbles of the B3 shoreline. This pause in lake level is marked in the Stockton Bar area by a boulder beach and is given the designation B6 (Fig. 9) (Burr and Currey, 1988). A major climate-driven subthreshold stage, termed the Keg Mountain Oscillation, is thought to have occurred following the formation of B6, resulting in net isostatic rebound of the basin floor. The lake returned once again to threshold control, and transgression continued at the rate of local hydrostatic subsidence to form the B7–B8 shorelines. From this viewpoint, the Stockton Bar extends westward ~2 km to South Mountain (Fig. 9). The concrete-like surface visible in the railroad cut is a tufa coating that was deposited during the youngest occupation of the Bonneville shoreline (Benson et al., 1990). Radiocarbon dating of the tufa, and a gastropod sample taken from the interstices of the tufa, gives ages of 14,730 ± 100 14C yr B.P. (Benson et al., 1990) and 14,420 ± 370 14C yr B.P. (Godsey et al., 2005), respectively. The Stockton spit extends southward from this viewpoint to the town of Stockton, Utah (Fig. 9). Past mining operations have exposed bedding planes on the eastern side of the spit. Aggregate mining operations are active on the southern tip of the spit. Directions to Stop 1.4 Proceed back downhill, turn left (west) at the first intersection with a well-defined dirt road. Stay left at the fork and follow the road around the top of the Stockton spit.

Don R. Currey memorial field trip Cumulative mi (km) 18.2

(29.3)

18.3

(29.5)

19.0 19.1

(30.6) (30.7)

19.5

(31.4)

20.1

(32.3)

20.4 20.6

(32.8) (33.2)

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Directions Carefully proceed downhill on the west side of the spit. Note the sand pit to the north. As recently as 2001, this pit contained highly detailed sedimentary structures including laminations, cross-beds, and ripple marks (Fig. 10). Off-road vehicle use has since destroyed all visible stratigraphy. A radiocarbon age of 14,730 ± 140 14C yr B.P. (Godsey et al., 2005) obtained on gastropods from this location suggests that this sand may be the winnowed sediments of B6, deposited shortly before the Bonneville flood. Turn south and proceed past the historic cemetery and the Stockton Jail. Turn right onto Sherman St. Turn right onto Silver Avenue and proceed west over railroad crossing. This highway is constructed on the Rush Valley equivalent of the Provo shoreline. Post-flood isolation of Rush Valley from the main body of Lake Bonneville caused the formation of a separate, smaller lake on the south side of the bar. At ~1539 masl (5050 fasl), this shoreline occurs at a much higher elevation than the Provo shoreline on the north side of the bar. Similar evidence of the Gilbert and Holocene highstands can be seen to the south of the highway. Turn right on dirt road marked “South Mountain Loop” by a rustic sign. We are driving over several transgressive-age shorelines as we head toward the Stockton Bar. Note the presence of several new homes on the shorelines to the east. Carefully proceed uphill (4-wheel drive may be necessary) and make a hard right at top. Turn left at fork. Proceed to hang-gliding launch site marked by white tufa outcrop and pink flagging. Stop 1.4: Stockton Bar (12, 383506E, 4480135N; NAD83).

Stop 1.4: Stockton Bar The apex of the Stockton Bar represents the B1 shoreline in this region (Fig. 9). From this perspective, you can see several transgressive-age shorelines and Provo-age beach ridges in Tooele Valley. Roughly 1 m below the B1 shoreline is the tufa-encrusted B8 shoreline. The vertical distance between the B8 shoreline and the Provo shoreline to the north emphasizes the magnitude of the Bonneville flood. The rusty-orange deposits visible to the north are residue of a tailings pond that was constructed in the lagoon behind the Bauer shoreline, a transgressive-age barrier (Fig. 9).

Figure 10. Don Currey, ca. 2000, standing in front of laminated sands at the apex of the Stockton Bar and the Stockton spit. A radiocarbon age of 14,730 ± 140 14C yr B.P. was obtained on gastropods from this location, indicating that these sands may be have been winnowed from the boulder beach at B6 and deposited just prior to the Bonneville flood. All of the stratigraphy visible in this photo has since been destroyed by off-road vehicle usage.

The Stockton Bar as a Geoantiquity In his final years, Don Currey was involved in research to document and preserve important landforms related to Lake Bonneville and Pleistocene Earth surface processes. He and his colleagues called these landforms “geoantiquities” and defined them as natural records of earth history that document environmental change on local, regional, and global scales (Chan et al., 2003) (Fig. 10). Sediments that make up Lake Bonneville landforms are typically well-rounded, well-sorted, and unconsolidated, making them prime aggregate material; therefore, these landforms are particularly susceptible to destruction or removal. The Stockton Bar tops the list of important geoantiquities, not only because of its scientific and educational merit, but also because of its historical, aesthetic, and recreational value. The bar has been a site of mining activity for at least 50 years, but excavation efforts have increased steadily with the growing aggregate needs of neighboring urban communities. In 2000, a request for a mining permit put the geoantiquities concept to the test. After many public speeches, field trips, community education campaigns, and partnerships with various conservation organizations, the struggle for preservation of the Stockton Bar has met with mixed success. A landfill and a tailings pond exist on the north side of the bar, homes have been developed on the transgressive shorelines to the south, and mining still threatens to remove deposits on the east side of the bar. However, new permits for mining have been stalled and a proposal to establish multi-use status for the Stockton Bar promises at least partial protection. Several of the locations we will visit on the remainder of this field trip are potential geoantiquity sites.

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428 Directions to Stop 1.5

Continue around the loop and back downhill to the paved highway. Turn left (east). Proceed through the town of Stockton to Connor Ave. (S.R. 36) and turn left (northeast) toward Tooele. Note the gravel foresets exposed in the B5 spit as we pass the gravel pit to the west. Turn left at the entrance to the Tooele Army

Depot. Turn right into the gravel lot just before the first building (Building 100) to get security passes. Proceed to guard checkpoint and reset mileage to zero. Cumulative mi (km) 0.0

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Figure 11. Surveyed topography of the Provo shoreline complex at three locations in the Bonneville Basin. P1, P3, P5, P7 and P9 are prominent beach ridges within the Provo shoreline complex. Dashed lines represent unsurveyed interpretations of topography. (A) Hogup Mountains; (B) Promontory Hollow; (C) Tooele Army Depot (after Godsey et al., 2005).

Directions From guard checkpoint, proceed NW on Main Ammo Road. The impressive ridges in the distance to the northwest are the Grantsville spits, formed during the transgression and highstand of Lake Bonneville. Turn left at the gate to the Tooele Army Depot Rifle and Pistol Range and follow the road into the gravel pit. Stop 1.5: P6–P7 Gravel Pit (12, 383413E, 4484360N; NAD83).

Stop 1.5: Tooele Army Depot (TEAD) Provo Shoreline, P6–P7 Gravel Pit The Provo shoreline began to form ca. 14.5 ka, following the Bonneville flood and the establishment of a new bedrock threshold at Red Rock Pass (Gilbert, 1890; Malde, 1968; Currey et al., 1984; Oviatt et al., 1992; Oviatt, 1997). The shoreline is essentially a ramp of prograding and aggrading beach ridges that are interrupted by abrupt downward steps (Fig. 11) (Burr and Currey, 1988). This morphostratigraphic signature may have resulted from landsliding in the outlet channel area and subsequent incision of landslide deposits that repeatedly raised and lowered the level of the threshold (Currey and Burr, 1988). However, changing climate conditions may also have affected the record of lake level in Provo-age deposits. Godsey et al. (2005) suggests that at least one subthreshold interval occurred during Provo time and that this interval may have been accompanied by isostatic rebound of the basin floor. Major Provo shorelines are designated P0–P9. Beach ridges or crests are designated with odd numbers: P1, P3, P5, P7 and P9 (Fig. 11). Topographic lows, or troughs at the base of each downward step, are given even number designations: P0, P2, P4, P6, and P8. P0 represents the point at which deposition began at the Provo shoreline. There are no designations for Provo subthreshold stages because the lake was originally thought to have remained open throughout the deposition of the Provo shoreline. Some of the best expressions of the Provo shoreline are located adjacent to and on the TEAD grounds. The Provo shoreline in the TEAD region consists of a series of ~80 cobble beach ridges that prograde laterally over 2300 m (Fig. 9). We will pass several of these beach ridges on the way to the gravel pit. Note the significant drop from P3 to P4 to the southwest. The P3–P4 drop is a signature feature of the Provo shoreline and is easily recognized throughout the Bonneville Basin (Fig. 11). The exposure in the gravel pit shows alternating coarse and fine material related to the P6–P7 beach ramp (Fig. 11). One interpretation of this outcrop is that the coarsening-up sequences represent annual or decadal depositional cycles.

Don R. Currey memorial field trip Directions to Stop 1.6

a significant downward oscillation, and subsequent isostatic rebound, occurred prior to its deposition (Fig. 11). Tufa deposits on P9 at TEAD and other locations in the basin suggest that this drawdown may have been related to a warming/drying event in the region that resulted in an increase of total dissolved solids in the lake waters (Godsey et al., 2002).

Return to Main Ammo road and turn right (southeast). Cumulative mi (km) 3.0

(4.8)

3.7

(5.6)

429

Directions Bear left on West Maintenance and Supply Road. Turn right onto gravel road, and proceed through gate to the gravel pit. Park at the eastern edge of the pit. Stop 1.6: P9 Spit (12, 384713E, 4485518N; NAD83).

Stop 1.6: Tooele Army Depot (TEAD) Provo Shoreline, P9 Spit This prominent spit has been identified as the P9 shoreline of the Provo shoreline complex (Godsey et al., 2005) (Figs. 9 and 11). Verifying the actual existence of a P9 component of the Provo shoreline has been somewhat problematic. The shoreline is not always readily identifiable, has few exposures, and may occur at a similar elevation as an older, transgressive-age shoreline (Sack, 1999). Radiocarbon dating on gastropods from this location produced an age of 13,580 ± 40 14C yr B.P. (Godsey et al., 2005). This age is inconsistent with the most recent version of the Bonneville hydrograph, which indicates that the lake remained at the Provo level from ca. 14.5 ka to 14 ka (Fig. 3). However, evidence from other locations supports the idea that the lake lingered at or near the Provo shoreline for a much longer period of time (Fig. 12) (Godsey et al., 2005). The considerably lower elevation of the P9 shoreline relative to the rest of the Provo shoreline complex may indicate that

Directions to Stop 1.7 Retrace route to Main Ammo Road and head southeast to exit the Depot. Turn left (NE) onto S.R. 36 and head toward Tooele. In Tooele, turn left (NW) onto 200 N. (S.R. 112) and continue to Grantsville. Turn left in Grantsville onto Main Street (S.R. 138) and follow road to intersection with I-80. Proceed through I-80 underpass to T-stop. Turn left (north) at T-stop to Stansbury Island and reset mileage to zero. Follow the road west and north through several bends and over two railroad crossings to the end of the pavement. Continue north over causeway to Stansbury Island. Algae, bacteria, and brine shrimp give the water a pink color in the evaporation ponds on either side of the causeway. Cumulative mi (km) 5.6

(9.0)

6.1

(9.8)

Directions Turn right (east) at crossroads (there is a sign indicating Broken Arrow Salt to the west) and drive to gravel pit. Floor of gravel pit. Shorelines are visible on the mountain facing us. Stop here if roads are wet or the route uphill is washed out. Otherwise, proceed carefully uphill to turn-around in marl outcrop.

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Figure 12. (A) Age and paleoelevation of radiocarbon samples detailed in Godsey et al. (2005), shown in relation to the Oviatt (1997) Lake Bonneville hydrograph. (B) Sample ages and paleoelevations shown with respect the Bonneville hydrograph from 15 ka to 11 ka. Dashed line indicates approximate lake level from Oviatt (1997) curve (after Godsey et al., 2005).

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(10.3)

Stop 1.7: Stansbury Gulch (12, 371852E, 4516581N; NAD83).

Stop 1.7: Stansbury Shoreline at Stansbury Gulch From Oviatt and Miller, 1997; Used with Permission The Stansbury oscillation is one of at least four major oscillations in lake level during the transgressive phase, each of which represents a significant change in water budget driven by climate change in the basin (Oviatt, 1997). For example, the Stansbury oscillation represented surface-area and water-volume changes of ~5000 km2 and 1000 km3, or relative changes of 18% and 50%, respectively (Oviatt et al., 1990). The other transgressivephase oscillations had similar magnitudes and represent climate changes probably associated with shifts in the mean position of storm tracks, which in turn were possibly determined by changes in the size and shape of the Laurentide ice sheet (Oviatt, 1997). We will examine exposures in Stansbury Gulch that show a section of the white marl and a wedge of tufa-cemented gravel that can be traced to the Stansbury shoreline. The exposures demonstrate that the Stansbury shoreline formed early in the lake history and that offshore stratigraphy can be linked to geomorphic features. Cream-colored sandstone outcrops form the west ridge of the short, steep valley, and gray limestone forms the east ridge. The tufa-cemented prominent shoreline high on the sandstone outcrops is the Provo shoreline, and the fainter shoreline about half way between the gravel pit (at the base of the mountain) and the Provo shoreline is the Stansbury shoreline. Both shorelines also can be seen on the limestone ridge. Most gravel exposed in the gravel pit at the base of the mountain was deposited during the initial transgression of Lake Bonneville. In some parts of the gravel pit, the white marl can be seen on the gravel near the top of the section; the marl is overlain by a few meters of cobbles, which were deposited during the rapid regression of the lake. The white marl was truncated in most places during this regression event. The stratigraphy and geomorphology of Stansbury Gulch have been described in several previously published guidebooks (see Currey et al., 1983; Green and Currey, 1988). Two or three thin sand beds in the diatomaceous marl in the lower parts of the gully exposures can be traced up-slope into thicker sand and then into a thick wedge of tufa-cemented gravel that is coincident with the Stansbury shoreline (Fig. 13). Radiocarbon ages of 20.7 ka determined on gastropods (Currey et al., 1983), and 23.3 ka on charcoal collected from the sand at the lower end of the gravel wedge, in addition to the stratigraphic relationships, indicate that the Stansbury shoreline formed during the transgressive phase of Lake Bonneville during an oscillation in lake level. Stratigraphic and geomorphic interpretations from other locations in the Bonneville Basin indicate that the total amplitude of the Stansbury oscillation was on the order of 45–50 m (150–165 ft) (Oviatt et al., 1990), although the evidence at Stansbury Gulch is insufficient in itself to demonstrate this.

FIELD TRIP DAY 2—BRIGHAM CITY DELTA, PUBLIC SHOOTING GROUNDS, AND HANSEL VALLEY Day 2 of the field trip involves an examination of the coarsegrained deposits of the Brigham City delta, the classic Gilbert shoreline deposits exposed at the Public Shooting Grounds, and an extensive section of late-transgressive, highstand, and early regressive deposits located in Hansel Valley. Directions to Stop 2.1 From Salt Lake City, travel north on I-15 toward Brigham City. Pass Exit 363 and exit into roadside rest area immediately after milepost 360. Stop 2.1: Overview of the Brigham City Delta (12, 412069E, 4591226N; NAD 83). Gilbert Deltas Gilbert’s 1890 study of Pleistocene Lake Bonneville was the first detailed geomorphic and stratigraphic study of gravely deltas. Gilbert is reported to have visited all of the lake’s deltas; however, he only discusses two locations in detail, the Bonneville-level delta at American Fork (Stop 3.5, Fig. 4) and the Provo-level Logan River delta (located in Cache Valley, northeastern Utah). From his observations of the lake’s coarse-grained deltas, Gilbert developed his topset-foreset-bottomset model (Fig. 14). However, recent gravel-pit exposures show that one of Gilbert’s original study localities, the Bonneville-level delta at American Fork, is composed almost entirely of subhorizontal gravel (topsets). Taking advantage of such gravel pit exposures not available to Gilbert, his model can be refined to show two end member deltas: (1) topset-dominated deltas deposited during the Bonneville transgression and highstand, and (2) foresetdominated deltas deposited at the Provo shoreline and during the Provo regression (Figs. 3 and 15; Table 1) (Milligan and Chan, 1998). Exposures at the Bonneville level of the Big Cottonwood Canyon and American Fork deltas display the horizontally stratified gravel that comprises the topset-dominated delta system (Stops 3.1 and 3.5, respectively; Fig. 4). Exposures at the Provo level of Big Cottonwood Canyon and Brigham City deltas display the steeply dipping gravel that comprises the foreset-dominated delta systems (Stops 3.1 and 2.2, respectively, Fig. 4). (Note: previous exposures of Provo level Big Cottonwood Canyon delta have been regraded and are now a golf course.) Three key factors contribute to the development and depositional styles of these Wasatch Front Gilbert deltas: active tectonism, rapid lake-level fluctuation, and drainage basin deglaciation. Slip on the Wasatch fault zone produced the steep drainage basins responsible for the overall coarse-grained nature of these Gilbert deltas. However, slip rates of 0.76–1.5 mm/yr, the average for the last 15 ka in study areas (Machette, 1988; Schwartz and Lund, 1988; Personius and Scott, 1992), are overprinted by lake-level change that can exceed 75 mm/yr (average for Bonneville level to Gilbert level regression; Fig. 3).

Don R. Currey memorial field trip

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Figure 13. Schematic measured sections from the walls of the gully at Stansbury Gulch (modified from Currey et al., 1983, and Green and Currey, 1988). The sections are simplified into three main units: lower marl (below the Stansbury sand and gravel); Stansbury sand and gravel, including the thick tufa-cemented gravel; upper marl (which represents deepwater deposition between the time of development of the Stansbury shoreline and the regression below the Provo shoreline); and Holocene colluvium and debris flows. Three radiocarbon ages have been obtained from these sections: 20,700 14C yr B.P. on Pyrgulopsis (Amnicola) shells and 23,300 14C yr B.P. on a small charcoal fragment, both from the Stansbury sand, and 24,900 14C yr B.P. on fine-grained CaCO3 from the lower marl (Green and Currey, 1988). This is Figure 10 from Oviatt and Miller, 1997; used with permission.

Figure 14. Gilbert’s classic topset-foreset-bottomset model of coarse-grained deltas (from Gilbert, 1890). A. Dip cross-“section of delta.” B. “Vertical section of a delta showing the typical succession of strata.”

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Figure 15. Generalized stratigraphic columns for gravely deltas at Brigham City, Big Cottonwood Canyon, and American Fork. Provo-level column at Big Cottonwood locality is enlarged at right to show detail. LST—low stand systems tract; TST—transgressive systems tract; HST— highstand systems tract; SB—sequence boundary; BS—bottomsets; FS—foresets; TS—topsets; DF—delta front. See Table 1 for lithology descriptions (from Milligan and Chan, 1998).

TABLE 1. LITHOFACIES DESCRIPTIONS FOR AMERICAN FORK, BIG COTTONWOOD, AND BRIGHAM CITY DELTAS Lithofacies

Description

Geometry and relations

Gravel topsets (TS)

Coarse pebble to cobble gravel, locally bouldery, clast to sandy matrix–supported, moderate to very poorly sorted, horizontal bedding.

Sheet geometry; limited localized channels at American Fork delta; overlies FS or DF.

Gravel foresets (FS)

Pebble to cobble gravel, clast supported, moderate to poorly sorted; some localities show interbedded silty sand, aquatic mollusks (Lymnaea/Amnicola), openwork gravel lenses, sand lenses, steeply dipping bedding (25–35°).

Sheet to wedge-like geometry; overlies and/or grades basinward to BS.

Sand and silt bottomsets (BS)

Very fine to fine sand and clayey silt, dropstones, aquatic mollusks (Lymnaea/ Amnicola), horizontal laminations, asymmetric and symmetric ripples, soft sediment deformation, low-angle to subhorizontal, interbedded gravel near FS contact.

Sheet geometry; overlain by FS, grades basinward to lake bottom deposits.

Delta front/beach fines (DF)

Coarse to very coarse sand with oblate pebbles and sandy clay, cross-bedding, wave ripples, discontinuous to continuous beds.

Sheet geometry; grades above and below to TS.

Note: Associated shoreline lithofacies not currently exposed in gravel pits visited on this trip include gravel barriers and back barrier fines.

Don R. Currey memorial field trip Drainage basin deglaciation and the release of glacial outwash played a role in sediment supply and, thus, the distribution of facies found at some localities. The effects of deglaciation are best seen at Big Cottonwood Canyon, where glaciers neared the Bonneville shoreline (Atwood, 1909; Hintze, 1988) producing topsets of subaerial glacial outwash. Evidence for glaciation (e.g., moraines and striations) is also found in the upper reaches of the American Fork drainage basin (Atwood, 1909; Hintze, 1988). Box Elder Canyon, the feeder canyon for the delta at Brigham City, shows no evidence of glaciation. Facies distributions were most strongly influenced by lakelevel fluctuation, which largely controls accommodation space and sediment supply. Topset-dominated deltas formed with increasing water depth created by climate-driven transgression to the Bonneville shoreline. Foreset-dominated deltas formed with decreasing water depth. Catastrophic lake-level drop due to the Bonneville flood and the subsequent climate-driven Provo regression not only greatly reduced accommodation space, but also provided abundant sediment supply by exposing unlithified Bonneville-level deltaic sediments for reworking. Stops 2.1 and 2.2 of this trip will examine the Provo-level foreset-dominated delta at Brigham City. Stops 3.4 and 3.5 will examine the Bonneville-level topset-dominated American Fork delta.

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Stop 2.1: Overview of the Foreset-Dominated Delta at Brigham City The objective of this stop is to observe the geomorphic setting of the foreset-dominated gravelly delta at Brigham City. Geologists not working with Quaternary deposits look at outcrops exposing internal stratigraphy and make inferences about region setting and landforms. With Quaternary deposits, we have the luxury of seeing the regional setting and geomorphology, but exposures of the internal stratigraphy in unlithified sediments can be limited in distribution or fleeting in duration. However, gravel pits, which maintain near-vertical, high walls offer an ever-changing glimpse of internal stratigraphy. To the east, a large “B” marks the wave-cut Bonneville shoreline on the mountain front above Brigham City. The Provo shoreline is seen ~100 m below. Extending from the canyon mouth at the Provo level is a gravelly delta exposed by large gravel pits (Fig. 16). From this perspective, the pit faces appear bright white due to the predominance of quartzite clasts derived from the Brigham Group up-canyon from the delta. No significant Bonneville transgressive deposits are found at this locality due to drainage basin physiography. Prior to deposition of the Provo-level foresets, Lake Bonneville extended roughly 5 km up an embayment into Box Elder Canyon (Fig. 17).

Bonneville Shorline Gravel pit exposures of Provo level delta

Figure 16. Distal view of Provo-level delta at Brigham City.

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H.S. Godsey et al. level. According to this scenario, fine-grained lake bottom or prodelta sediments (deposited while the coarse-grained sediments were stored in the embayment) should underlie the steeply dipping foresets. However, ground-penetrating radar shows another 32 m of coarse-grained foresets beneath those exposed (Smith and Jol, 1992), indicating that any prodelta sediments are well below gravel-pit exposures. Directions to Stop 2.2 Exit the rest area and resume traveling north on I-15. Reset mileage to zero. Cumulative mi (km)

Figure 17. Schematic diagrams (in sequential time slices) of Gilbert delta at Brigham City. (A) At the Bonneville highstand (when lake water extended up an embayment), deposition was often limited to the embayment and narrow canyon. (B) During the Provo-lake level (when lake water met the canyon mouth), recently exposed Bonneville-level deposits provided abundant sediment to build deltas. (C) Post-Provo erosion then incised the Provo-level delta.

Fluvial gravels were deposited and stored within this embayment until the catastrophic lowering of base level by ~100 m due to the Bonneville flood. This drop resulted in plentiful sediment supply but reduced accommodation space. Box Elder Creek began to rework the unconsolidated sediments (confined in the prior embayment) and redeposited them in the abundant progradational foresets (lowstand systems tract) now found at the Provo

1.0 1.7

(1.6) (2.7)

1.3 1.0

(2.1) (1.6)

Directions Take Exit 364 toward Brigham City. In Brigham City, turn left onto Main Street (2nd signal). Turn right on 200 South (3rd signal). Turn left at Staker and Parson Sand and Gravel Pit (north of S.R. 90 at the mouth of Box Elder Canyon). Check in at pit office. Stop 2.2: Staker and Parson Sand and Gravel Pit (12, 417054E, 4595115N; NAD83). HARD HATS ARE REQUIRED TO ENTER PIT AREA.

Stop 2.2: Steeply Dipping Gravel Foresets (Lowstand Systems Tract) at Staker and Parson Sand and Gravel Pit At this stop, pit walls expose well-developed, steeply dipping foresets, deposited as the lowstand systems tract during the Provo regression (Fig. 18). These steeply dipping (25° to 35°) foresets are characterized by clast-supported gravel in a matrix of poorly sorted fine- to very coarse-grained sand, with localized open-framework (i.e., no matrix) gravel lenses. The moderate to poorly sorted clasts range from pebbles to cobbles and show no preferred orientation or grading. The steep, generally westward dip (away from the canyon mouth sources) of the beds suggests

Figure 18. Well-developed, steeply dipping (~25° to 30°) gravel foresets of the Provo-level delta at Brigham City (Stop 2.2). Located in the north side of the delta, this pit face strikes NW (left) and is ~70 m high.

Don R. Currey memorial field trip

435 112°20’

deposition in deltaic foresets due to mass transport processes (gravity flow) on the slipface. Collectively, foreset beds exhibit a sheet to wedge geometry.

112°15’ m fro w g flo prin am lt S tre Sa

Blue Spring Hills

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Directions to Stop 2.3 Return to Main Street. Turn left onto Main Street (Hwy 13) and travel north. Bear left, staying on Hwy 13 toward Corinne. Bear left onto Hwy 83. Reset mileage to zero. Cumulative mi (km) (2.6)

2.6 10.3

(4.2) (16.6)

11.8

(19.0)

13.8

(22.2)

The highway drops to the modern flood plain of the Bear River. The Bear River probably brought the largest water and clastic sediment influx to Lake Bonneville. East of the Bear River, flats produced during the Gilbert regression merge with a broad, low-relief delta plain. Bear left on Hwy 83. Smelly hot springs are common near the road as we drive along the base of Little Mountain. The Stansbury, Provo, and Bonneville shorelines are prominently displayed to the northeast on Little Mountain. Stop 2.3 and 2.4: Public Shooting Grounds (12, 390135 E, 4606755 N, NAD83, and 390412 E, 4607007 N, NAD83).

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g Stockin Molly's

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LEGEND

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Figure 19. Map of the Public Shooting Grounds area. Gray irregular areas represent platforms; intervening white areas are mudflats.

Public Shooting Grounds At the Public Shooting Grounds (a wildlife area managed by the Utah Division of Wildlife Resources), deposits of the late regressive phase of Lake Bonneville are overlain by lacustrine, wetland, and eolian deposits, and all are exposed along the margins of flat- to round-topped topographic platforms that stand several meters above the surrounding mudflats. The shapes of some of the platforms are subequant in planview, whereas others are elongate and in some places radiate from springwater source areas (Fig. 19). The margins of the platforms provide exposures of the late Pleistocene and early Holocene deposits in this area. Don Currey recognized the importance of the exposures at the Public Shooting Grounds for understanding the age of the Gilbert shoreline and related deposits. He summarized his interpretations in 1990 (Currey, 1990; Fig. 15), and a simplified version of his summary figure is presented here (Fig. 20). Currey interpreted the Public Shooting Grounds sequence as representing channel sands, coastal marsh deposits, and fluviodeltaic sand deposited during the Gilbert paleolake cycle, which he interpreted as lasting ~2000 radiocarbon years and culminating at ca. 10.3 ka (Benson et al., 1992; Fig. 5). The Public Shooting Grounds exposures were important to Currey in the development of his ideas about the “pre-Gilbert red beds,” which he thought of as the product of the reworking and oxidizing of fine-grained

4 km

41°30’

Figure 20. Simplified version of Currey’s (1990, Figure 15) interpretation of the stratigraphic sequence at the Public Shooting Grounds (this is Figure 26 of Oviatt and Miller, 1997, used here with permission from BYU Geology Studies). Currey (1990) reported radiocarbon ages of gastropods from the channel sands of 10,920, 10,990, 11,570, and 11,990 14C yr B.P.

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sediment of Lake Bonneville in the lake as it regressed below the Provo shoreline. One of Currey’s major contributions through his work at the Public Shooting Grounds was to demonstrate that the lake had not risen above the altitude of the Public Shooting Grounds exposures after the formation of the Gilbert shoreline (also noted by Miller et al., 1980). At this location we will examine post–Lake Bonneville deposits that have been useful for testing the age of the Gilbert shoreline (first dated directly by Currey) and for understanding the Holocene history of wetlands at the margin of Great Salt Lake. We have obtained new radiocarbon ages on organic materials from the Public Shooting Grounds and have interpreted the stratigraphic sequence as indicating that the deposits here represent the regression of Lake Bonneville, the transgression to the Gilbert shoreline, and subsequent deposition in spring-fed wetlands and in eolian and/or sheetwash environments (Oviatt et al., 2005). We will visit two sites at this stop, one of which is located next to Utah Hwy 83 (section 5 of Oviatt et al., 2005); the other will require a short (0.3 km) hike north of the highway (section 6, Fig. 19). Stop 2.3: Section 5 The road cut at section 5 (road cut on the south side of Hwy 83, ~4 mi west of Little Mountain) (Oviatt et al., 2005) provides a view of a typical stratigraphic section at the Public Shooting Grounds. The lower part of the cut exposes finegrained sediments of Lake Bonneville (unit 1, Fig. 21), which are overlain by ripple-laminated fine sand (unit 2), sandy mud (unit 3W), and massive, poorly sorted fine sand (unit 4). Unit 1 is pale reddish brown mud in its lower part that grades upward

Figure 21. Photo of a typical exposure at the Public Shooting Grounds (PSG) (along the eastern side of the platform, SE of section 5, Fig. 18 herein). Stratigraphic-unit boundaries are readily apparent in the field because of color differences. Stratigraphic units are labeled; 1r—unit 1 (reddish brown); 1g—unit 1 (greenish gray).

into pale greenish-gray mud, which in places contains sand-filled mud cracks. We cored unit 1 at section 5 using a Livingston corer and found it to be >4 m thick and to contain the Hansel Valley basaltic ash (Miller et al., 1995) and an ostracode assemblage and faunal sequence typical of Lake Bonneville deposits elsewhere in the basin (the “White Marl” of Gilbert, 1890). Unit 2 ripplelaminated sand can be traced to within a few vertical meters of the Gilbert shoreline on the south flank of Little Mountain, and we interpret it as representing shallow offshore deposition during the transgression of Great Salt Lake to that shoreline. Unit 2 is stratigraphically overlain by mud, muddy sand, and sand, some of which is rich in organics, that we interpret as having been deposited in wetland environments (unit 3W; Fig. 21). Mollusk shells are common in places in this stratigraphic unit. In places channels cut through units 2 and 1, and are filled with sand, mud, abundant mollusk shells, and organics (unit 3C) (Fig. 22). At most exposures in the Public Shooting Grounds area, wetland deposits are overlain by massive, poorly sorted sand that we interpret as having been deposited by wind and reworked by sheet-flow processes (unit 4). Sand of unit 4 forms low dunes along the margins of some platforms and underlies the flat surfaces of the platforms. Stop 2.4: Section 6 At section 6 (0.3 km north of Hwy 83, NE of section 5 along the east side of a topographic platform), units 1, 2, 3W, and 4 and the sandy fill of a wetland channel (3C) are exposed along the eastern side of a platform (Fig. 22). Here the side of the channel can be seen where it cuts through unit 2 and into unit 1; the channel-fill sediments (3C) are overlain by organic-rich mud of unit 3W, and sand of unit 4 caps the section. At this site we will discuss possible interpretations of the origin of the sandy 3C channel fill. Calibrated radiocarbon age ranges of organic material from this channel overlap with calibrated age ranges of organics from the base and middle of unit 2 (see Figure 12 in Oviatt et al., 2005), but as shown by the cross-cutting stratigraphic relationships at section 6, unit 3C is younger than unit 2. This suggests that soon after the deposition of unit 2 (and the formation of the Gilbert shoreline), base level (lake level?) dropped rapidly, inducing incision into the slightly older Lake Bonneville deposits, then the channel filled rapidly with sandy sediment carrying abundant mollusks. Radiocarbon Results We have obtained a total of 17 new radiocarbon age estimates from the Public Shooting Grounds and have compared these results to those previously reported by Miller et al. (1980), Murchison (1989), Currey (1990), and Light (1996) (see Oviatt et al., 2005, for details). There is a large degree of scatter in the radiocarbon ages of mollusks, and the mollusk ages are consistently older than the ages of carbonized plant fragments from the same stratigraphic units. Therefore, we have concluded that a variable hard-water effect is apparent in the mollusk ages (due

Don R. Currey memorial field trip to the spring water) and that the ages of fragments of emergent aquatic plants are superior and provide a reliable and reproducible chronology. We partially tested this hypothesis by dating the shell of a snail that had died within a year of when we collected it and obtained an age of ca. 600 14C yr B.P. Note that in unit 3C at section 6 the mollusk radiocarbon ages range from 11,970– 14,550 14C yr B.P. and they are significantly older than the two radiocarbon ages of carbonized plant fragments (9850 and 9980 14 C yr B.P.) (Fig. 22).

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The ages and depositional environments of the Public Shooting Grounds stratigraphic units are summarized in Table 2. See Oviatt et al. (2005) for stratigraphic and geochronologic details. Summary of Late Pleistocene and Holocene Events at the Public Shooting Grounds The Public Shooting Grounds stratigraphy indicates the following sequence of events during the late Pleistocene and early Holocene:

section 6 altitude 1293 m unit 4

unit 3W 1m

unit 3C

unit 2

14.55 ± 60 12.79 ± 60 11.97 ± 40 14.36 ± 50 9.98 ± 40 9.85 ± 40

unit 1

Figure 22. Measured section and photograph of section 6. Radiocarbon ages of samples collected from unit 3C are shown with triangles (open triangles represent mollusk shells and black triangles represent carbonized plant fragments). Modified from Figure 10 of Oviatt et al. (2005).

TABLE 2. STRATIGRAPHIC UNITS IN THE PUBLIC SHOOTING GROUNDS AREA Unit

Age*

4

~6.6† ~7.5 7.8–6.9 8.7–7.6 9.7–9.5 11.1–10.5 10–9.8 11.9–11.2 10.5–10 12.9–11.1 ~28–11# ~31–13

3M§ 3W 3C 2 1

Lithology Poorly sorted fine sand and silt, massive to weakly bedded, in sheets and dune forms. Muddy sand; massive, bioturbated, poorly sorted; locally includes thinly bedded clean sand. Poorly sorted mud and sandy mud; locally organic rich; bioturbated; carbonate nodules. Mollusk-rich sand and muddy sand.

Interpretation Eolian, sheetwash deposition Wetlands, mud flats Wetlands Channel fills

Ripple-laminated medium to fine sand; thin mud drapes; Shallow lacustrine, Gilbert sand includes well-rounded gravel near Little Mountain. Pale reddish brown mud, grading downward to brown mud, Lacustrine, Bonneville mud, oxidized then dark gray mud; upper part locally pale greenish gray; and cracked in upper part ostracodes.

Note: modified from Table 2 of Oviatt et al. (2005). *Estimated radiocarbon age (14C ka) range, followed by calibrated age range (cal ka) in italics. † Based on luminescence age (see Oviatt et al., 2005). § This unit is well developed in the Molly’s Stocking area (Figure 18) south of sections 5 and 6. # From Oviatt et al. (1992).

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1. Regression of Lake Bonneville and deposition of muddy sediment (unit 1) prior to ca. 11,000 14C yr B.P.; 2. Lake lowering to below the altitude of Public Shooting Grounds exposures, oxidation of Bonneville muds, and formation of mud cracks; 3. Lake transgression to the Gilbert shoreline and deposition of unit 2 sometime between 10,500 and 10,000 14 C yr B.P.; 4. Lake lowering, channel incision, and channel filling in a short period immediately after 10,000 14C yr B.P.; 5. Continued wetland deposition between 10,000 14C yr B.P. and at least 7000 14C yr B.P. (local wetland deposition probably continued throughout the Holocene); and 6. Deposition of eolian and/or sheetwash sand beginning in the mid-Holocene.

18.2

(29.3)

19.6

(31.5)

23.3

(37.5)

23.4

(37.7)

28.1 29.2

(45.2) (47.0)

31.1

(50.1)

32.0

(51.5)

34.4 35.6

(55.4) (57.3)

Directions to Stop 2.5 Reset vehicle mileage to zero and proceed west on Hwy 83 across marshland of the Public Shooting Grounds. Cumulative mi (km) 6.1

(9.8)

8.1

(13.0)

11.3

(18.2)

12.8

16.2

(20.6)

(26.1)

Directions Lampo Junction. Turn left toward Golden Spike National Historic Site. We cross a broad alluvial plain built on Bonneville marl that is truncated to the south, just below this altitude, by Gilbert transgressive lake deposits. Junction; bear right. Proceed up Promontory Mountains, crossing original transcontinental railroad beds. Transgressive-phase spits of Lake Bonneville are prominent landforms to the south, at the north end of the Promontory Mountains. The spits lie above the Provo shoreline, which is close the altitude of the highway here, and therefore formed as transgressive lake features. During the catastrophic drop of lake level from Bonneville to Provo shorelines (Bonneville flood), little geomorphic work could be performed. Intersection with gravel road; continue straight on the gravel road. Golden Spike National Monument, erected to commemorate the historic meeting of the transcontinental (Union and Central Pacific) railroads, is to the left. We are driving along an unconformity cut into Miocene tuff during Pliocene time. Alluvial sediment on the tuff, but beneath Lake Bonneville sediment, yielded Pliocene fossils and volcanic ash (Nelson and Miller, 1990). We have descended from the crest of the North Promontory Mountains to the Bonne-

ville shoreline, which to the south forms a conspicuous barrier beach. Cross double Provo barrier beach. As in most places where the Provo shoreline is well formed, it consists of two beaches ~3 m different in altitude; here they are widely spaced due to the gentle slope. Gilbert (1890) noted the double character of erosional segments of the Provo shoreline, but offered no explanation. Currey (Currey and Burr, 1988) notes three or four steps in depositional Provoshoreline segments at a number of locations around the basin and suggests that landsliding and scour in the overflow threshold at Red Rock Pass, Idaho, complicated by ongoing isostatic rebound, controlled lake level throughout the basin during the development of the Provo shoreline. The Provo shoreline is expressed as wave-cut notches on both sides of the road. Notches were cut into Miocene tuffaceous sediment of the Salt Lake Group. Continue straight. Route following the old railroad grade is to the left. Double Gilbert barrier beach. Quarry pit on the left is in one of the beaches. Descend to and cross the mud flats. Turn right on gravel road toward northeast. Climb onto Gilbert barrier beach from Gilbert erosional notch. Gilbert barrier beach is visible to the south of road. Its crest is slightly above 4260 ft. To the east, it merges with the level of the road. Cross a degraded scarp created by the 1934 Hansel Valley earthquake and previous faulting events. The Hansel Valley fault has displaced the surface of the Gilbert spit down to the east. The fault strikes slightly east of north, and its scarp is visible north of the road for some distance. In the mud flats south of the road, the mud was cracked and mud volcanoes formed in 1934, but only scattered evidence for the location of the fault can be seen now. Bear left. Stop 2.5: Hansel Valley Wash (12, 362400 E, 4626100 N; NAD83).

Hansel Valley Wash Modified from Oviatt and Miller, 1997; Used with Permission Hansel Valley Wash contains a marl section that is notable for several features, including: (1) its lateral continuity for several km along Hansel Valley, (2) presence of the Hansel Valley

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439

basaltic ash near the base, and (3) soft-sediment disruption of the marl, possibly induced by seismicity. Stop 2.5 Hansel Valley Wash Stop along the main road next to an obscure road on the right in a greasewood plain. Take care not to drive in the greasewood; it destroys tires! Walk ~1.6 km (1 mi) along this obscure road and continue into Hansel Valley Wash as the road ends. The first kilometer of the wash has been modified by bulldozer, but eventually the wash turns to its original northeasterly orientation. Proceed up this original wash ~1 km until the marl section is ~3 m thick as exposed in walls on the east side of the gully. The Bonneville section at this stop is fairly complete because it was deposited on a low-gradient valley floor at a relatively low elevation (1320 m; 4330 ft). Here we can observe the sequence of facies changes in the marl that can be seen at many similar sections around the basin. Coarse-grained deposits at the base of the section are interpreted as marking the initial transgression of Lake Bonneville (Fig. 23). The coarse sand grades upward into blocky mud that contains oxidized root holes; we interpret this unit as having been deposited in a marsh or lagoon environment at the margin of the transgressing lake. Overlying the transgressive deposits is a sequence of laminated marl 1 m (3.2 ft) thick (early-transgressive and Stansbury), which grades upward into more massive, greenish gray to pink marl ~1.3 m (4 ft) thick (deepwater marl). The upper contact of the massive marl is abrupt, and the overlying bed of ripple-laminated sand and sandy marl is ~12 cm (0.4 ft) thick (the Bonneville flood bed). Its upper contact is gradational into another massive marl (Provo marl), which coarsens upward and is disrupted in its upper part by modern soil development. Ostracodes and diatoms (Fig. 23) support the interpretation of this sequence as a cycle, representing the transgression, deep water, and regression of Lake Bonneville. One thing to speculate on at this section is the origin of the ripple-laminated beds in the Bonneville flood bed (the 12-cmthick bed between the two massive marls). Our interpretation is that during and immediately after the Bonneville flood, when lake level dropped catastrophically by ~100 m, vast areas of finegrained deepwater sediments would have been stranded between the Bonneville and Provo shorelines. That sediment would have begun washing into the lake immediately after the flood, perhaps as debris flows and landslides that would have created turbidity currents on the lake bottom. In Hansel Valley, slumping of finegrained sediments both above and below lake level might have been enhanced by earthquake activity. Two centimeters above the base of the laminated marl is a thin (1 cm) bed of brown basaltic ash, which we have named the Hansel Valley ash (Miller et al., 1995). We have found the Hansel Valley ash at many localities in northern Utah, including in the Burmester core at the south end of Great Salt Lake (Oviatt et al., 1999). At all localities where the Hansel Valley ash has been found, including sediment cores from Great Salt Lake (Spencer et al., 1984), the ash bed is within a few centimeters of the base

Figure 23. Photo of Hansel Valley Wash marl section. T—transgressive mud and sand; LM—laminated (transgressive) marl; DWM—deepwater marl; BF—Bonneville flood bed; PM—Provo marl. Ostracode samples: W—Cyl, Ce, Ca, Lc, Lsa; V—Ce, Cc, Ca, Lc; U—Cc, Ce, Ca, Lc, Lsa; T—Lc, Cc, Ca, Ce; S—Lc, Ls, Ca, Cc; R—Lc, Ca?; Q—Lc, Ca, Cc; P—Ca, Lc; O—Ca, Lc; N—Ca, Lc; M—Lc, Ca; L—Lc, C sp.; K—Lc, Cc, Ca, Cd; J—Lc, Cc, Ca, Cd; I—Lc, Cc; H—Lc, Cc, Ls; G—Cc, Ls; F—Cc, Ls; E—Ls, Cc; D—Ls, C sp.; C—Ls; B—Ls; A—no ostracodes. Ostracode abbreviations: Ca—Candona adunca; Cc—Candona caudata; Ce—Candona eriensis; Cyl—Cytherissa lacustris; Lc—Limnocythere ceriotuberosa; Ls—Limnocythere staplini; Lsa—Limnocythere sappaensis. Diatoms from Hansel Valley Wash section identified by Platt Bradbury, May 27, 1992, in samples W, V, and C–G: W—Cyclotella ocellata (cold open water); V—Synedra acus (fresh open water), S. ulna, Cyclotella ocellata, C. caspia??, Fragilaria brevistriata, F. leptostauron; C–G—Fragilaria brevistriata (shallow, moderately saline water), F. construens v. subsalina, Epithemia, Mastogloia, Navicula, Amphora, Surirella, Pinnularia, others. Modified from Figure 25 in Oviatt and Miller, 1997; used with permission.

of the Bonneville section. A radiocarbon age of 26,500 14C yr B.P. for a sample collected near the ash in a Great Salt Lake sediment core (Thompson et al., 1990) is the best available age for the eruption. Exposures in a tributary gully to Hansel Wash (referred to as West Gully) in Hansel Valley suggest that Lake

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440

Bonneville was close to an elevation of 1335 m (4380 ft) at the time the Hansel Valley ash was erupted. Despite extensive field efforts, we have not yet identified the source vent of the ash, but its chemistry is similar to that of basalts ~20 km to the west in the Curlew Valley area (Miller et al., 1995). Note the common disrupted beds containing small faults and folds below and including the Hansel Valley ash. These features may have been caused by nearby small seismic events or larger distant events. Upstream several km, convoluted beds and hummocky cross-stratification are common in the section beneath the deepwater beds that lie below the flood bed. Robison and McCalpin (1987) suggested that these features indicate several local earthquakes, some of which displaced parts of the marl section in West Gully.

lacustrine sediments deposited during the Bonneville highstand. The trip ends at the American Fork delta, a topset-dominated gravelly delta that was deposited during the Bonneville transgression and highstand. Day 3 Road Log From Salt Lake City, travel south on I-15, and exit onto I-80 east. Proceed east along I-80 ~4.8 mi, then exit onto I-215 south. Proceed along I-215 ~6.5 mi; exit onto 6200 South. Turn east and reset mileage to zero. Cumulative mi (km) 0.0

(0.0)

1.6 1.8

(2.6) (2.9)

Directions to Stop 2.6 Return to gravel road, reset mileage to zero, and turn south. Cumulative mi (km) 7.5

(12.1)

11.3

(18.2)

Directions Bear straight on gravel road toward the southwest. Road to left returns to Golden Spike and Brigham City. Stop 2.6 Monument Point (12, 362400 E, 4626100N; NAD83). Drive south toward Lone Rock, and stop near the southern tip of the wave-cut bluffs.

Directions 6200 South heads east then turns south and becomes Wasatch Blvd. The gravel pit on the left of Wasatch Blvd. exposes horizontally stratified glacial outwash composed of poorly sorted, clast-supported, pebbles and cobbles deposited as delta topsets. Wasatch Blvd. is at the elevation of the Provo shoreline, the highest bench at the top of the gravel pit is the Bonneville shoreline. Turn right at Fort Union Blvd. Pull out on shoulder of road on right. Stop 3.1: Big Cottonwood Canyon Delta (12, 432920E, 4496855N; NAD83).

Big Cottonwood Canyon Stop 2.6: Monument Point This stop provides another opportunity to examine an exposure of the Lake Bonneville marl. This section was well exposed in the late 1980s and early 1990s after Great Salt Lake waves had undercut and freshly eroded the bluff during the high lake levels of the mid-1980s. The Hansel Valley ash is exposed here where it overlies sand at the base of the section. Look for places where the ash bed is draped over quartzite dropstones. The slightly increasing trend in the percentage of total inorganic carbon (% TIC) (Fig. 24) above the Hansel Valley ash is similar to the trend seen in cores of Bonneville sediments from the bottom of Great Salt Lake (Oviatt et al., 2005, Fig. 8). The stratigraphic position of the Bonneville flood is difficult to place in this section. FIELD TRIP DAY 3—BIG COTTONWOOD CANYON, LITTLE COTTONWOOD AND BELLS CANYONS, AMERICAN FORK DELTA Day 3 of the trip begins at the mouth of Big Cottonwood Canyon where the remains of a classic Gilbert-style delta have been exposed by down-cutting during the regression of Lake Bonneville. The trip continues south on Wasatch Boulevard to the mouth of Little Cottonwood and Bells Canyons to examine glacial sediments and landforms and their temporal relation to

Contribution of New Material and Current Research by Elliott Lips Sediments at the mouth of Big Cottonwood Canyon were deposited in fluvial and deltaic environments during the last two deep lake cycles of the Bonneville Basin (Scott, 1988a). Milligan and Chan (1998) identified and described distinct architectural elements of Gilbert-type gravely deltas exposed in the gravel pits at the mouth of Big Cottonwood Canyon. The majority of the sediments preserved between the Bonneville and Provo shorelines consist of horizontally stratified glacial outwash composed of poorly sorted, clast-supported, pebbles and cobbles deposited as delta topsets. The sediments observed below the Provo shoreline consist of steeply dipping pebble to cobble delta foresets, interfingering with horizontally bedded, fine sand and silt delta bottomsets. The present topography of the area results from incision of the delta by Big Cottonwood Creek. Stop 3.1: Big Cottonwood Canyon Delta and the Geomorphic Response of Big Cottonwood Creek to the Regression of Lake Bonneville In this part of Salt Lake Valley, the Bonneville shoreline has isostatically rebounded to ~1582 masl (5190 fasl) and the

Don R. Currey memorial field trip Provo shoreline to ~1481 masl (4860 fasl). Stop 3.1 is at ~1484 masl (4870 fasl), essentially at the Provo shoreline. Wasatch Boulevard is built on the Provo shoreline, and the gravel pit to the northeast exposes the coarse-grained delta topsets deposited between the Bonneville and Provo shorelines. The Bonneville highstand is marked to the north by the flat-topped surface with towers for a gun club; to the south, the shoreline is out of view but is close to the highest ridge in our view with houses. Delta foresets and bottomsets were previously exposed to the northwest and west, below the elevation of the Provo shoreline. Mining, urbanization, and construction of a golf course have since removed these exposures. This stop provides an opportunity to examine the geomorphic response of Big Cottonwood Creek to changing lake levels during the regression of Lake Bonneville. During the Bonneville transgression and highstand (16 ka to 14.5 ka), gravel was deposited at the mouth of the canyon in either broad braided streams or fan-deltas as the stream entered the lake. Throughout this ~1500 yr episode, the creek migrated back and forth across the delta top, depositing sediments radially out from the canyon mouth. The current edge of the Bonneville-level delta is located ~1000 m to the west. Changing base level during the Bonneville flood (ca. 14.5 ka) caused the creek to incise into the recently deposited deltaic sediments. Given that the sediments were saturated and unconsolidated, it is likely that down-cutting kept pace with the lowering lake level. Thus, at the end of the Bonneville flood, the creek incision would likely have been near the elevation of the Provo shoreline. In cross section, the incision would likely have been narrow and v-shaped because the stream would not have had time to migrate laterally and erode into the adjacent sediments. Following the Bonneville flood, deposition of the delta continued at the Provo shoreline. Because base level stayed more or less constant, the creek did not cut down farther into the Bonneville-level delta; however, over the course of ~800 yr, the creek would have likely migrated laterally, widening the valley incised into the delta. While the lake was at the Provo shoreline, the creek would likely have been a shallow, braided stream carrying large amounts of glacial outwash and reworked Bonneville-level delta sediments. The Provo-level delta would have been contiguous between our location and the Provo shoreline to the north along Wasatch Boulevard. The Provo-level delta front is located ~2500 m west of Stop 3.1. In response to climate change, Lake Bonneville began to slowly recede from the Provo shoreline until ca. 11.5 ka when it reached the level of modern Great Salt Lake (Godsey et al., 2005) (Fig. 12). Big Cottonwood Creek responded to this continual lowering of base level by down cutting and lateral erosion. As the lake receded, the creek cut down through the Provo-level delta, deepening the valley between our stop and Wasatch Boulevard. Once again, these sediments were reworked and transported basinward. However, because the lake did not stabilize for any significant period of time during its regression, there is no post Provo-aged delta recognized in the Salt Lake Valley. Instead,

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Figure 24. Lab data from a sediment core taken at the Monument Point exposure (modified from Figure 4 in Oviatt et al., 1994). Boundaries of lithologic units are approximate. TIC—total inorganic carbon.

the deposits consist of stream alluvium of all sizes graded to the receding lake (Personius and Scott, 1992). The valley below Stop 3.1 represents the floodplain of modern day Big Cottonwood Creek, which is graded to its confluence with the Jordan River, its base level control. Directions to Stop 3.2 Return to Wasatch Boulevard, turn right (south) and reset mileage to zero. Cumulative mi (km) 2.2

(3.6)

3.3 3.4

(5.4) (5.5)

Directions Veer right at fork in order to stay on Wasatch Blvd. Turn right on Little Cottonwood Rd. Park on the shoulder on the right side of the road.

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(5.5)

Stop 3.2: G.K. Gilbert Geologic Interpretive Park (12, 432415E, 4491751N; NAD83). Walk along the sidewalk 200 ft to the east.

Little Cottonwood and Bells Canyons Contribution of New Material and Current Research by Elliott Lips Little Cottonwood and Bells Canyons offer a unique opportunity to examine the temporal relation between lake and glacier fluctuations and their paleoclimate implications. The convergence point of these canyons is one of only two locations in the western United States where Pleistocene glaciers extended below the shorelines of pluvial lakes (the other is Mono Basin in the eastern Sierra Nevada). Beginning with G.K. Gilbert (1890), one fundamental question that has been addressed is the relationship between the timing of the glaciers and the lake. Scott (1988b) provides a summary of different interpretations of the age of both glacial deposits at the canyon mouth and the highstand deposits of Lake Bonneville. Don Currey contributed significantly to understanding this relationship based on his exhaustive knowledge of the radiocarbon chronology of Lake Bonneville and his recognition of the stratigraphic and geomorphic relations between the lake and the moraines. Even more significantly, Madsen and Currey (1979) provided the first and, as of 2005, only radiocarbon date constraining the age of the moraines. Madsen and Currey demonstrated two ages of glaciation beyond the canyon mouths, Bull Lake and Pinedale. They named the associated tills Dry Creek Till and Bells Canyon Till, respectively. We will examine the type localities for these tills and the sites visited by Don Currey in his work on the glacial and lacustrine chronologies. In addition, we will examine the results of recent investigations (Lips et al., 2005) that provide a different age of the glaciation and an explanation for discrepancies with previous interpretations. Stop 3.2: G.K. Gilbert Geologic Interpretive Park— Overview of Late Pleistocene Glaciation The view from Salt Lake County’s G.K. Gilbert Geologic Interpretive Park provides textbook examples of erosional and depositional glacial features. Little Cottonwood Canyon displays a U-shaped cross section, characteristic of glacially eroded bedrock valleys (Fig. 25). Canyon walls expose Tertiary white quartz monzonite of the Little Cottonwood stock (Crittenden et al., 1973). Extending beyond the canyon mouth to the south is a prominent left lateral moraine from Little Cottonwood Canyon. This moraine has been mapped as Bells Canyon age, which is equivalent to till of Pinedale age (correlative to marine oxygen isotope stage [MIS] 2; Shackleton and Opdyke [1973]) mapped elsewhere in the Rocky Mountains (Personius and Scott, 1992). Gilbert (1890) mapped four distinct left lateral moraines extending beyond the mouth of Little Cottonwood Canyon (Fig. 25). On the north side of the canyon, the right lateral moraine is less

voluminous but can be identified by the position of the large quartz monzonite boulders on the hillside and extending past the canyon mouth. Recent mapping (Lips et al., 2005) has identified at least three Bells Canyon–aged right lateral moraines extending past the canyon mouth (Fig. 26). The left lateral moraine contains more till because ten north-facing cirques supplied ice and debris to the main west-flowing valley and piedmont glacier in Little Cottonwood Canyon (Madsen and Currey, 1979). The elevation of the G.K. Gilbert Geologic Interpretive Park is 1586 masl (5203 fasl), near the elevation of the Bonneville highstand. However, at this location, we are on the hanging wall of the Wasatch Fault and the Bonneville highstand sediments have been displaced downward to an elevation of ~1561 masl (5120 fasl). The left lateral moraines of Little Cottonwood Canyon are offset by multiple down-to-the-west traces of the Wasatch Fault. Normal faulting, in conjunction with antithetic faults, has created grabens upslope of the terminal moraine at Bells Canyon and across the right lateral moraines of Little Cottonwood Canyon (Fig. 26). The sediments lying between the two lateral moraines consist of poorly sorted pebbles and cobbles in a matrix of sand and silt deposited as stream alluvium and glacial outwash during the regressive phase of Lake Bonneville (Personius and Scott, 1992). Meltwater streams that deposited the outwash also eroded and/or buried end moraines that would have existed past the canyon mouth. In addition, there is no evidence of shorezone features from the Bonneville highstand preserved in the outwash deposits. Directions to Stop 3.3 Return to Little Cottonwood Rd., heading west, and reset mileage to zero. Cumulative mi (km) 0.4 0.9

(0.7) (1.5)

1.0

(1.6)

Directions Turn left on 3100 East. 3100 East turns to the right and becomes Mt. Jordan Road (10000 South). Turn left into the parking lot for Dimple Dell Park. Stop 3.3: Inspiration Point (12, 431670E, 4491094W; NAD83). Follow the narrow path just to the right of the wooden sign 0.17 mi to Inspiration Point. The relatively flat surface that extends south and west at the beginning of the walk is the top of the delta created from Little Cottonwood Creek and Dry Creek (Bells Canyon) at the Bonneville highstand. The delta surface is covered by up to 10 ft of post-Bonneville eolian sand.

Stop 3.3: Inspiration Point—Temporal Relation between Till of Bells Canyon Age and the Bonneville Highstand The prominent landforms to the east are the well-preserved left lateral, terminal, and right lateral moraines extending past the

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Figure 25. Oblique aerial view of Little Cottonwood Canyon (left) and Bells Canyon (right) looking to the east. Field trip stops 3.2 and 3.3 are shown with white stars. The white dotted lines delineate the crests of four left lateral moraines that extended beyond the mouth of Little Cottonwood Canyon. The large flat surface in the lower center of the photograph is the delta created from sediments likely derived from both Little Cottonwood and Bells Canyons during the Bonneville highstand. In the right-center of the photograph are the well-preserved left lateral, terminal, and right lateral moraines of Bells Canyon. Approximate location of the Wasatch Fault shown by white dashed line; note the fault cuts across and displaces the moraines.

Figure 26. Oblique aerial view of the mouth of Little Cottonwood Canyon looking toward the southeast. Field trip Stop 3.2 is shown with the white star in the right of the photograph. The white dotted lines delineate the crests of the two youngest right lateral moraines that extend beyond the canyon mouth and the most prominent left lateral moraine. The right lateral moraines are discontinuous across a graben of the Wasatch Fault. The white dashed lines delineate the upper limit of the main fault and the lower limit of the corresponding antithetic fault. The large flat surface in the lower left corner with the nursery plots is the delta surface created at the mouth of Little Cottonwood Canyon during the Bonneville highstand. Trenches at the contact of the penultimate moraine and the Bonneville-age delta are shown with short black lines. The left lateral moraine from Little Cottonwood Canyon and the Bells Canyon moraines are in the upper right of the photograph.

mouth of Bells Canyon (Fig. 25). The relatively flat surface to the southwest and northwest is the top of the Bonneville-age delta that likely received sediments from Little Cottonwood Canyon and Bells Canyon (Lips et al., 2005). The valley between our stop and the end moraine of Bells Canyon contains Dry Creek, which follows Dimple Dell Road in its upper reach. Directly across Dry Creek from our stop is the location where Madsen and Currey (1979) described the stratigraphy of the tills, and nearby is the site where they collected a sample for radiocarbon dating.

The lowest exposed unit is the Dry Creek till, which contains numerous clasts of quartz monzonite, virtually all of which disintegrate to grus under light impact (Madsen and Currey, 1979). Based on physical characteristics, stratigraphic position, and correlations by previous workers, Madsen and Currey (1979) interpreted the Dry Creek till to be associated with marine oxygen isotope stage 6 (Shackleton and Opdyke, 1973) glaciation and correlative with deposits of Bull Lake age elsewhere in the Rocky Mountains. Madsen and Currey (1979) obtained a date

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of 26,080 ± 1200/1100 14C yr B.P. on total organic carbon from the Majestic Canyon soil, a mature paleosol developed on the Dry Creek till. Overlying the Majestic Canyon soil is till of Bells Canyon age, identified on the basis of fewer weathered quartz monzonite clasts and well preserved geomorphic features. In a few isolated locations, a thin wedge of lacustrine sediments associated with the Bonneville highstand overlies the Bells Canyon till (Personius and Scott, 1992). Based on this stratigraphic sequence, the age of the Majestic Canyon soil, which provides a maximum limiting age for the upper till, and the assumed age of the Bonneville sediments at the time, Madsen and Currey (1979) concluded that the most probable age of the maximum extent of the Bells Canyon till was in the range of 19 ka to 20 ka. However, they did note that prominent moraines of Bells Canyon age existed ~1 km upstream from the maximum limit and may be up to a few thousand years younger than the maximum. Madsen and Currey (1979) contributed to the understanding of the temporal relation between the glaciers and Lake Bonneville in three ways. First, they provided the only (as of 2005) radiocarbon date that has been utilized in constraining the age of the till. Their date clearly demonstrated that the youngest till was correlative with MIS2 and Pinedale-age glaciation. This was significantly different than the interpretation by Richmond (1964) that all the till beyond the canyon mouth was of Bull Lake age. Second, Madsen and Currey (1979) recognized that the Bonneville sediments on the Bells Canyon till were only found in discontinuous and isolated patches. Third, they recognized multiple moraines of Bells Canyon–aged till that extended beyond the canyon mouth. To determine the temporal relation between the glaciers and Lake Bonneville, most previous workers relied on the stratigraphic position of the till and lacustrine sediments exposed at the surface (Scott, 1988b). However, slope wash and/or eolian sediment largely covers this contact. In addition, urbanization and road construction have eliminated most surface exposures in the vicinity of Bells Canyon. Recent investigations (Lips et al., 2005) have examined the stratigraphic relations exposed in newly created road cuts and in trenches excavated specifically to expose the contact between the lacustrine and glacial sediments. In addition, examination of the geomorphic relation between glacial and lacustrine landforms provides information on their temporal relation (Lips et al., 2005). Finally, cosmogenic 10Be dating techniques provide the exposure age of quartz monzonite boulders on the crests of moraines at Little Cottonwood Canyon (Lips et al., 2005). Bonneville sediments exposed at the surface immediately below Wasatch Boulevard were previously mapped by Personius and Scott (1992) but have been removed and/or buried by construction of Wasatch Blvd. Recent excavation at The Boulders at Bells Canyon subdivision exposed Bonneville highstand sediments and till from Bells Canyon. The Bonneville sediments were observed overlying and, in turn, overlain by till recognized as Bells Canyon age by Madsen and Currey (1979) and Personius and Scott (1992). Lips et al. (2005) interpret these tills to be two different advances during MIS2, one occurring before the highstand of Lake Bonneville, and one after the highstand. Two advances explains

why there are only isolated patches of the Bonneville sediments found on the till; the younger advance covered most, but not all, of the lacustrine sediments. Evidence for at least one glacial advance after the highstand of Lake Bonneville is also present in trenches excavated at the toe of the penultimate right lateral moraine beyond the mouth of Little Cottonwood Canyon (Fig. 26). At two locations, till of Bells Canyon age is observed overlying lacustrine sediments at the elevation of the Bonneville highstand, clearly indicating that there was at least one advance after the lake reached the Bonneville highstand (Lips et al., 2005). Geomorphic evidence at and beyond the mouths of Bells and Little Cottonwood Canyons also suggests at least one glacial advance after the Bonneville highstand (Lips et al., 2005). Deposition of the large delta surface that extends west and southwest from the canyons required a sustained period of stream discharge. However, a left lateral moraine presently forms a topographic divide that would block sediment discharge from Little Cottonwood Canyon to the delta. In addition, the well-preserved end moraine of Bells Canyon (note the lack of a significant breach) would similarly block sediment from Bells Canyon (Fig. 25). Therefore, a delta of the size and location observed must have formed before the moraines. A similar relation exists for the large delta surface that extends northwest from Little Cottonwood Canyon. The two youngest right lateral moraines could not have existed prior to the Bonneville highstand because they would have prevented sediment from forming the northern portion of the delta (Fig. 26). Additional geomorphic evidence is found in comparing the size of the valley that has been eroded by Dry Creek to the notch eroded in the end moraine of Bells Canyon (Fig. 25). The crosssectional area of the valley of Dry Creek is several orders of magnitude larger than the present cut in the moraine. This suggest that the valley was eroded by water flows of a much greater discharge and that subsequent to the valley erosion, the moraines advanced past the canyon mouth to their present configuration (Lips et al., 2005). If this sequence is correct, there was at least one advance of the glaciers past the canyon mouth after the Bonneville highstand because lacustrine sediments of Bonneville age are cut by the stream erosion. Finally, the fluvial terraces located in the floor of Dry Creek are graded to Provo deltas and shorelines (Machette and Currey, 1988). This suggests that there was sustained flow through Dry Creek after the Bonneville flood while the lake was at the Provo stillstand. Cosmogenic 10Be exposure ages of moraine boulders at Little Cottonwood Canyon provide additional evidence of the timing of glaciation (Lips et al., 2005). 10Be concentration is proportional to the time the boulder has been on the moraine surface and thus determines the date of the moraine formation (Gosse and Phillips, 2001). Seven large quartz monzonite boulders on the crests of the two youngest (based on geomorphic position) right lateral moraines were sampled (Fig. 26). Only the largest boulders were selected from the crest. To avoid boulders subjected to spalling or erosion, only those with minimal pitting were selected. The mean exposure age of the seven boulders was 15.9 ± 0.7 10Be ka (range between 15.2 ± 0.4 10Be ka and 16.9 ±

Don R. Currey memorial field trip 0.4 10Be ka) (Lips et al., 2005). To compare moraines ages with Lake Bonneville chronology, the 10Be years were converted to 14 C years with CALIB 5.0.1 (Stuiver et al., 2005). Assuming the 10 Be time scale is nearly equivalent to a calendrical time scale, the moraine age is ca. 13.4 14C ka. This age for the last advance of glaciers is in agreement with the stratigraphic and geomorphic evidence at Little Cottonwood and Bells Canyons. In addition, it suggests that climate conditions were cold and/or wet as late as 13.4 14C ka, approximately when the lake was at the Provo level and spilling through Red Rock Pass (Godsey et al., 2005). Directions to Stop 3.4 Return to Mount Jordan Road heading west and turn left onto Little Cottonwood Road (9600 South). Drive west on Little Cottonwood Road (it becomes 9400 South) to the onramp for I-15 South. From I-15, take Exit 287 onto Hwy 92 east, toward Alpine/Highland. Reset mileage to zero. Cumulative mi (km) 2.2

(3.5)

2.8

(4.5)

Directions Turn left on Center Street (this dirt road is a public right of way). Continue up the dirt road to Stop 3.4: Overview of American Fork Delta (12, 427939E, 4476933N; NAD83).

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transgressive–highstand system tract (Fig. 28). These gravel topset deposits consist of horizontal clast-supported pebble and cobble gravel with lenses of silty sand deposited during the Bonneville transgression and highstand. The sandy matrix contains less than 3% clay. Gravel beds (m-scale thickness) are distinguished by grain size variations, but internally show no obvious grading. This topset facies has a sheet-like geometry with an intervening 9 m section of delta front–beach fines (topset–delta front–topset sequence). The coarse-grained topsets are generally inferred to represent bedload transport in planar sheets under high-energy (sediment gravity) flow conditions. Some workers present various flow processes and depositional mechanics for subaerial and subaqueous horizontally stratified, fan delta gravels (e.g., Nemec et al., 1984). In these Bonneville deposits, the topsets probably represent a range of mass flow processes. The intervening 9 m section of delta front–beach fines is characterized by sandy clay and coarse-grained to very coarse– grained sand with granules and oblate pebbles. Sedimentary structures include wave ripples and tabular cross-bedding. The cross-bedding suggests a southerly flow direction (parallel to the shoreline) and is likely to have been created by littoral currents. The presence of oblate pebbles and symmetric ripples

Stop 3.4: Overview of the Topset-Dominated American Fork Delta Looking east, the gravelly American Fork delta is clearly seen at the Bonneville shoreline. In map view, the Bonnevillelevel American Fork delta is geomorphically expressed as a classic fan shape or delta “Δ” shape, incised by the modern American Fork stream channel (Fig. 26). However, in contrast to alluvial fans that have concave-up long profiles that steepen toward the mountain front, the Bonneville-level American Fork delta is subhorizontal near the mountain front then shows a sharp break in slope farther toward the basin (Fig. 27). Directions to Stop 3.5 Return to Hwy 92 and continue east for 7.4 km (4.6 mi) to Westroc Inc., Highland Pit (north side of S.R. 92 at the mouth of American Fork Canyon). Turn left into pit and check in at scales. Stop 3.5: Westroc Inc., Highland Pit (12, 435433E, 4476026N; NAD83). HARD HATS ARE REQUIRED TO ENTER PIT AREA. Stop 3.5: Horizontally Stratified Gravel Topsets (Transgressive– Highstand Systems Tract) at Westroc Inc. Highland Pit. The Bonneville-level American Fork delta seen at this stop exemplifies the topset-dominated system deposited as the

Figure 27. Aerial view of Bonneville-level American Fork delta ca. 1970. Note incision by the modern American Fork river channel.

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Figure 28. Extensive Bonneville-level gravel pit exposure of gravely topset (TS), sandy delta front (DF), and gravely topset sequence (TS) of the American Fork Delta. This amphitheater view is approximately east-west at the left and north-south at the right of photo. Pit face is ~55 m high

suggest shallow water, wave-influenced deposition in a delta front–beach environment. The occurrence of this fine-grained beach facies amidst coarse-grained delta topsets may be attributed to a major downward oscillation (Machette, 1988). The drop in lake level during this oscillation (Fig. 3) probably caused the American Fork River to incise a channel through the delta, thus transferring the river deposition westward into the basin. This river channel (until filled) would have cut off the coarse-grained sediment supply (during the oscillation rise and final transgression to the Bonneville shoreline), allowing the accumulation of the finer-grained beach and delta-front sands. ACKNOWLEDGMENTS We are grateful to our colleagues who worked with us on many aspects of the research included in this field guide, including Jack McGeehin, Cecile Zachary, Shannon Mahan, Rick Forester, David Madsen, Joe Rosenbaum, Dave Bedford, Stephanie Dudash, Chris Hoglund, Josh Howard, Marjorie Chan, Alisa Felton, Ian Schofield, and Shizuo Nishizawa. Scott Ritter, editor of Brigham Young University Geology Studies, kindly granted us permission to reprint parts of our GSA fieldtrip guidebook from 1997 (Oviatt and Miller, 1997). Parts of this work were supported by National Science Foundation grants SBR-9817777 (to M. Chan and D. Currey), EAR-9809241 (to P. Jewell, D. Currey, and M. Chan), and ESI0329669 (to M. Chan). REFERENCES CITED Atwood, G., 2002, Storm-related flooding hazards, coastal processes, and shoreline evidence of Great Salt Lake, in Gwynn, J.W., ed., Great Salt Lake: An overview of change: Salt Lake City, Utah Department of Natural Resources, p. 43–53. Atwood, G., 2004, Tribute to Donald R. Currey, January 24, 1934–June 6, 2004: Friends of Great Salt Lake Newsletter, Summer 2004, v. 10, no. 4, p. 16–17. Atwood, W.W., 1909, Glaciation of the Uinta and Wasatch Mountains: U.S. Geological Survey Professional Paper 61, 96 p.

Balch, D.P., Cohen, A.S., Schnurrenberger, D.W., Haskell, B.J., Valero Garces, B.L., Beck, J.W., Cheng, H., and Edwards, R.L., 2005, Ecosystem and paleohydrological response to Quaternary climate change in the Bonneville Basin, Utah: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 221, p. 99–122, doi: 10.1016/j.palaeo.2005.01.013. Benson, L.V., Currey, D.R., Dorn, R.I., Lajoie, K.R., Oviatt, C.G., Robinson, S.W., Smith, G.I., and Stine, S., 1990, Chronology of expansion and contraction of four Great Basin lake systems during the past 35,000 years: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 78, p. 241–286, doi: 10.1016/0031-0182(90)90217-U. Benson, L.V., Currey, D.R., Lao, Y., and Hostetler, S., 1992, Lake-size variations in the Lahontan and Bonneville Basins between 13,000 and 9000 14 C yr B.P.: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 95, p. 19–32, doi: 10.1016/0031-0182(92)90162-X. Burr, T.N., and Currey, D.R., 1988, The Stockton Bar, in Machette M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 66–73. Chan, M.A., Currey, D.R., Dion, A.N., and Godsey, H.S., 2003, Geoantiquities in the urban landscape, in Heiken, G., Fakundiny, R., and Sutter, J., eds., Earth science in the cities: A reader: American Geophysical Union Monograph, p. 21–42. Crittenden, M.D., Jr., Stuckless, J.S., Kister, R.W., and Stern, T.W., 1973, Radiometric dating of intrusive rocks in the Cottonwood area, Utah: U.S: Geological Survey Journal of Research, v. 1, no. 2, p. 173–178. Currey, D.R., 1980, Coastal geomorphology of Great Salt Lake and vicinity, in Gwynn, J.W., ed., Great Salt Lake: A scientific, historical and economic overview: Utah Geological and Mineral Survey Bulletin 116, p. 69–82. Currey, D.R., 1990, Quaternary palaeolakes in the evolution of semidesert basins, with special emphasis on Lake Bonneville and the Great Basin, U.S.A.: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 76, p. 189–214, doi: 10.1016/0031-0182(90)90113-L. Currey, D.R., and Burr, T.N., 1988, Linear model of threshold-controlled shorelines of Lake Bonneville, in Machette M.N., ed., In the footsteps of G.K. Gilbert— Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 104–110. Currey, D.R., and Oviatt, C.G., 1985, Durations, average rates, and probable causes of Lake Bonneville expansion, still-stands, and contractions during the last deep-lake cycle, 32,000 to 10,000 yrs ago, in Kay, P.A., and Diaz, H.F., eds., Problems of and prospects for predicting Great Salt Lake levels—Proceedings of a National Oceanic and Atmospheric Administration Conference, March 26–28, 1985: Center for Public Affairs and Administration, University of Utah, Salt Lake City, Utah, p. 9–24. Currey, D.R., Oviatt, C.G., and Plyler, G.B., 1983, Lake Bonneville stratigraphy, geomorphology, and isostatic deformation in west-central Utah: Utah Geological and Mineral Survey Special Studies, v. 62, p. 63–82.

Don R. Currey memorial field trip Currey, D.R., Atwood, G., and Mabey, D.R., 1984, Major levels of Great Salt Lake and Lake Bonneville: Utah Geological and Mineral Survey Map 73, scale 1:750,000. Doelling, H.H., Willis, G.C., Jensen, M.E., Hecker, S., Case, W.F., and Hand, J.S., 1990, Geologic map of Antelope Island, Davis County, Utah, Map 127: Salt Lake City, Utah Geological Survey, plate 2, scale 1:24,000. Gilbert, G.K., 1890, Lake Bonneville: U.S. Geological Survey Monograph 1, 438 p. Gilluly, J., 1929, Possible desert-basin integration in Utah: Journal of Geology, v. 36, p. 672–682. Godsey, H.S., Currey, D.R., Felton, A.K., and Chan, M.A., 2002, Refining the record of Pleistocene lake level change, Lake Bonneville, Utah: Evidence of climate-driven oscillations from the Provo shorezone: Geological Society of America Abstracts with Programs, v. 34, no. 6, p. 368. Godsey, H.S., Currey, D.R., and Chan, M.A., 2005, New evidence for an extended occupation of the Provo shoreline and implications for regional climate change, Pleistocene Lake Bonneville, Utah: Quaternary Research, v. 63, p. 212–223, doi: 10.1016/j.yqres.2005.01.002. Gosse, J.C., and Phillips, F.M., 2001, Terrestrial in situ cosmogenic nuclides: Theory and application: Quaternary Science Reviews, v. 20, p. 1475– 1560, doi: 10.1016/S0277-3791(00)00171-2. Green, S.A., and Currey, D.R., 1988, The Stansbury shoreline and other transgressive deposits of the Bonneville lake cycle: Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 55–57. Hintze, L.F., 1988, Geologic History of Utah: Department of Geology, Brigham Young University, 202 p. Hunt, C.B., 1982, Pleistocene Lake Bonneville, ancestral Great Salt Lake, as described in the notebooks of G.K. Gilbert, 1875–1880: Brigham Young University Geology Studies, v. 29, 225 p. Hunt, C.B., Varnes, H.D., and Thomas, H.E., 1953, Lake Bonneville-Geology of northern Utah Valley, Utah: U.S. Geological Survey Professional Paper 257-A, 99 p. Lemons, D.R., Milligan, M.R., and Chan, M.A., 1996, Paleoclimatic implications of late Pleistocene sediment yield rates for the Bonneville Basin, northern Utah: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 123, p. 147–159, doi: 10.1016/0031-0182(95)00117-4. Light, A., 1996, Amino acid paleotemperature reconstruction and radiocarbon shoreline chronology of the Lake Bonneville Basin, USA [M.S. thesis]: Boulder, University of Colorado, 142 p. Lips, E.W., Marchetti, D.W., and Gosse, J.C., 2005, Revised chronology of late Pleistocene glaciers, Wasatch Mountains, Utah: Geological Society of America Abstracts with Programs, v. 37, no. 7 (in press). Machette, M.N., 1988, American Fork Canyon, Utah: Holocene faulting, the Bonneville fan-delta complex, and evidence for the Keg Mountain oscillation, in Machette, M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 89–96. Machette, M.N., and Currey, D.R., 1988, Road log from Salt Lake City to the northern part of Utah Valley and return, in Machette M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 75–77. Machette, M.N., and Scott, W.E., 1988, A brief review of research on lake cycles and neotectonics of the eastern Basin and Range province, in Machette M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 7–14. Madsen, D.B., and Currey, D.R., 1979, Late Quaternary glacial and vegetation changes, Little Cottonwood Canyon area, Wasatch Mountains, Utah: Quaternary Research, v. 12, p. 254–270, doi: 10.1016/00335894(79)90061-9. Malde, H.H., 1968, The catastrophic late Pleistocene Bonneville Flood in the Snake River Plain, Idaho: U.S. Geological Survey Professional Paper 596, 52 p. McCoy, W.D., 1987, Quaternary aminostratigraphy of the Bonneville Basin, western United States: Geological Society of America Bulletin, v. 98, no. 1, p. 99–112, doi: 10.1130/0016-7606(1987)98<99: QAOTBB>2.0.CO;2. Miller, D.M., Nakata, J.K., Oviatt, C.G., Nash, W.P., and Fiesinger, D.W., 1995, Pliocene and Quaternary volcanism in the northern Great Salt Lake area

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and inferred volcanic hazards: Utah Geological Association Publication, v. 24, p. 469–482. Miller, R.D., Van Horn, R., Scott, W.E., and Forester, R.M., 1980, Radiocarbon date supports concept of continuous low levels of Lake Bonneville since 11,000 yr B.P.: Geological Society of America Abstracts with Programs, v. 12, p. 297–298. Milligan, M.R., and Chan, M.A., 1998, Coarse-grained Gilbert deltas: facies, sequence stratigraphy and relationships to Pleistocene climate at the eastern margin of Lake Bonneville, northern Utah, in Shanley, K.W., and McCabe, P.J., eds., Relative role of eustasy, climate, and tectonism in continental rocks: Society for Sedimentary Geology (SEPM) Special Publication no. 59, p. 177–189. Morrison, R.B., 1965, Quaternary geology of the Great Basin, in Wright, H.E., and Frey, D.G., eds., The Quaternary of the United States: Princeton, New Jersey, Princeton University Press, p. 265–285. Murchison, S.B., 1989, Fluctuation history of Great Salt Lake, Utah, during the last 13,000 years [Ph.D. thesis]: Salt Lake City, University of Utah, 137 p. Nelson, M.G., and Miller, D.M., 1990, A Tertiary record of the giant marmot Paenemarmota sawrockensis in northern Utah: Contributions to Geology, v. 28, p. 31–37. Nemec, W., Steel, R.J., Porebski, S.J., and Spinnanger, A., 1984, Domba Conglomerate, Devonian, Norway: Process and lateral variability in a mass flow-dominated lacustrine fan-delta, in Koster, E.H., and Steel, R.J., eds., Sedimentology of gravels and conglomerates: Calgary, Canadian Society of Petroleum Geologists Memoir 10, p. 295–320. O’Connor, J.E., 1993, Hydrology, hydraulics, and geomorphology of the Bonneville flood: Geological Society of America Special Paper 274, 83 p. Oviatt, C.G., 1997, Lake Bonneville fluctuations and global climate change: Geology, v. 25, p. 155–158, doi: 10.1130/0091-7613(1997)025<0155: LBFAGC>2.3.CO;2. Oviatt, C.G., and Miller, D.M., 1997, New explorations along the northern shores of Lake Bonneville, in Link, P.K., and Kowallis, B.J., eds., Mesozoic to recent geology of Utah: Brigham Young Geology Studies v. 42, pt. II, p. 345–371. Oviatt, C.G., Currey, D.R., and Miller, D.M., 1990, Age and paleoclimatic significance of the Stansbury shoreline of Lake Bonneville, northeastern Great Basin: Quaternary Research, v. 33, p. 291–305. Oviatt, C.G., Currey, D.R., and Sack, D., 1992, Radiocarbon chronology of Lake Bonneville, eastern Great Basin, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 99, p. 225–241, doi: 10.1016/00310182(92)90017-Y. Oviatt, C.G., Habiger, G., and Hay, J., 1994, Variation in the composition of Lake Bonneville marl: A potential key to lake-level fluctuations and paleoclimate: Journal of Paleolimnology, v. 11, p. 19–30, doi: 10.1007/ BF00683268. Oviatt, C.G., McCoy, W.D., and Reider, R.G., 1987, Evidence for a shallow early or middle Wisconsin lake in the Bonneville Basin, Utah: Quaternary Research, v. 27, p. 248–262, doi: 10.1016/0033-5894(87)90081-0. Oviatt, C.G., Thompson, R.S., Kaufman, D.S., Bright, J., and Forester, R.M., 1999, Reinterpretation of the Burmester core, Bonneville Basin, Utah: Quaternary Research, v. 52, p. 180–184, doi: 10.1006/qres.1999.2058. Oviatt, C.G., Miller, D.M., McGeehin, J.P., Zachary, C., and Mahan, S., 2005, The Younger Dryas phase of Great Salt Lake, Utah, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 219, no. 3-4, p. 263–284, doi: 10.1016/j.palaeo.2004.12.029. Personius, S.F., and Scott, W.E., 1992, Surficial geologic map of the Salt Lake City segment and parts of adjacent segments of the Wasatch Fault Zone, Davis, Salt Lake, and Utah Counties, Utah: U.S. Geological Survey Miscellaneous Investigations Series Map I-2106, scale 1:50,000. Richmond, G.M., 1964, Glaciation of the Little Cottonwood and Bells Canyons, Wasatch Mountains, Utah: U.S. Geological Survey Professional Paper 454-D, 41 p. Robinson, R.M., and McCalpin, J.P., 1987, Surficial geology of Hansel Valley, Box Elder County, Utah, in Kopp, R.S., and Cohenour, R.E., eds., Cenozoic geology of western Utah: Utah Geological Association Publication 16, p. 335–349. Sack, D.S., 1999, The composite nature of the Provo level of Lake Bonneville, Great Basin: North America: Quaternary Research, v. 52, p. 316–327. Schofield, I., Jewell, P.W., Chan, M.A., Currey, D.R., and Gregory, M., 2004, Shoreline development, longshore transport, and surface wave dynamics, Pleistocene Lake Bonneville, Utah: Earth Surface Processes and Landforms, v. 29, p. 1675–1690, doi: 10.1002/esp.1121.

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Schwartz, D.P., and Lund, W.R., 1988, Paleoseismicity and Earthquake recurrence at Little Cottonwood canyon, Wasatch fault zone, Utah, in Machette, M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 82–85. Scott, W.E., 1988a, Deposits of the last two deep-lake cycles at Point of the Mountain, Utah, in Machette M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 86–88. Scott, W.E., 1988b, Temporal relations of lacustrine and glacial events at Little Cottonwood and Bells Canyon, Utah, in Machette M.N., ed., In the footsteps of G.K. Gilbert—Lake Bonneville and neotectonics of the eastern Basin and Range province: Geological Society of America Annual Meeting Field Trip Guidebook, Utah Geological and Mineral Survey Miscellaneous Publication 88-1, p. 78–81. Scott, W.E., McCoy, W.D., Shroba, R.R., and Rubin, M., 1983, Reinterpretation of the exposed record of the last two cycles of Lake Bonneville, west-

ern United States: Quaternary Research, v. 20, no. 3, p. 261–285, doi: 10.1016/0033-5894(83)90013-3. Shackleton, N.J., and Opdyke, N.D., 1973, Oxygen isotope and paleomagnetic stratigraphy of equatorial Pacific core V23-238: oxygen isotope temperatures and ice volumes on a 105 year and 106 year scale: Quaternary Research, v. 3, p. 39–55, doi: 10.1016/0033-5894(73)90052-5. Smith, D.G., and Jol, H.M., 1992, Ground-penetrating radar investigation of a Lake Bonneville delta, Provo level, Brigham City, Utah: Geology, v. 20, p. 1083–1086, doi: 10.1130/0091-7613(1992)020<1083: GPRIOA>2.3.CO;2. Spencer, R.J., Baedecker, M.J., Eugster, H.P., Forester, R.M., Goldhaber, M.B., Jones, B.F., Kelts, K., McKenzie, J., Madsen, D.B., Rettig, S.L., Rubin, M., and Bowser, C.J., 1984, Great Salt Lake and precursors, Utah: The last 30,000 years: Contributions to Mineralogy and Petrology, v. 86, p. 321–334. Stuiver, M., Reimer, P.J., and Reimer, R., 2005, CALIB radiocarbon calibrations: http://radiocarbon.pa.qub.ac.uk/calib/ (Accessed 10 Mar. 2005). Thompson, R.S., Toolin, L.J., Forester, R.M., and Spencer, R.J., 1990, Accelerator-mass spectrometer (AMS) radiocarbon dating of Pleistocene lake sediments in the Great Basin: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 78, p. 301–313, doi: 10.1016/0031-0182(90)90219-W.

Printed in the USA

Geological Society of America Field Guide 6 2005

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment Lee Amoroso U.S. Geological Survey, 2255 North Gemini Drive, Flagstaff, Arizona 86001, USA Jason Raucci Department of Geology, Northern Arizona University, Flagstaff, Arizona 86011, USA

ABSTRACT The Hurricane Fault is one of the longest and most active late Cenozoic normal faults in southwestern Utah and northwestern Arizona. This fault shows evidence of tectonic activity during the late Tertiary and Quaternary, neotectonism involving the Hurricane Fault as well as the Toroweap Fault imply encroaching Basin and Range extension onto the Colorado Plateau. Paleoseismology investigations suggest that the Hurricane Fault poses a seismic hazard to the southwestern Utah area. During the trip, we will examine evidence of late Pleistocene and earliest Holocene(?) surface-rupturing faulting along the Shivwits and Whitmore Canyon sections of the fault. The Hurricane Fault separates the Uinkaret and Shivwits plateaus and displacement along the fault produced the spectacular Hurricane Escarpment. We will see late Quaternary landforms related to back-wasting and mass movement along the Hurricane Escarpment and look at evidence of the style and age estimates of late Pleistocene fan deposition. Keywords: Hurricane Fault, paleoseismology, neotectonism, alluvial fan, colluvium. INTRODUCTION The Hurricane Fault (Fig. 1) is the longest and most active of the late Cenozoic down-to-the-west normal faults in southwestern Utah and northwestern Arizona. The Hurricane Fault crosses the Arizona Strip between the Utah border and Grand Canyon in close proximity to St. George, Utah (Fig. 1). Although the Arizona portion of the Hurricane Fault crosses sparsely populated terrain, much of populous southwestern Utah lies within 75 km of the Shivwits section. Two significant, historic seismic events have occurred in the region. An ~M6 earthquake occurred in the Pine Valley, Utah, area in 1902 (Williams and Tapper, 1953). A M5.8 earthquake in the St. George area in 1992 caused minor structural damage in southwestern Utah, triggered a large landslide near the entrance to Zion National Park 45 km from the epi-

center (Christensen, 1995), and caused numerous rockfalls along the Hurricane cliffs (G.H. Billingsley, 2000, personal commun.). Several recent paleoseismic investigations have addressed the potential for larger earthquakes than those of the historic record. These workers have suggested that the threshold magnitude for surface rupture along faults within the Intermountain Seismic Belt (ISB; Fig. 1) in Utah is 6 < M < 6.5 (Arabasz et al., 1992; Doser, 1985; Smith and Arabasz, 1991). Fault scarps and other evidence of Quaternary faulting suggest that there is potential for M > 7 earthquakes along the Hurricane Fault (Stewart et al., 1997; Stenner and Pearthree, 1999; Amoroso et al., 2004). This field trip guide introduces evidence for late Quaternary ruptures on the Hurricane Fault in Arizona, considers the neotectonics implications, and places the late Quaternary deformation within the context of the encroachment of Basin-and-Range style

Amoroso, L., and Raucci, J., 2005, Paleoseismology and geomorphology of the Hurricane Fault and Escarpment, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 449–477, doi: 10.1130/2005.fld006(20). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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deformation onto the Colorado Plateau. The deposits that result from geomorphic and tectonic processes along the Hurricane Escarpment are also introduced.

border. The abrupt decrease in displacement, from more than 2500 m in Utah to 250–400 m on the Arizona portion of the fault may be the result of differences in crustal properties between the Basin and Range and Colorado Plateau (Stewart et al., 1997).

GEOLOGIC OVERVIEW Stratigraphy The Hurricane Fault is a 250-km-long, high-angle, normal fault extending from near Cedar City, Utah, to south of Grand Canyon (Fig. 1). In southwestern Utah, from Cedar City to the Arizona border, the Hurricane Fault zone is the physiographic boundary between the Colorado Plateau and Basin and Range (Figs. 1 and 2) (Arabasz and Julander, 1986). In northwestern Arizona, the Grand Wash Fault, 50 km west of the Hurricane Fault, forms the physiographic boundary (Anderson and Mehnert, 1979; Mayer, 1985). Ongoing east-west–oriented extension in the Basin and Range–Colorado Plateau boundary zone in northwestern Arizona and southwestern Utah is accommodated along long, down-to-the-west, normal fault zones, including the Hurricane, Toroweap-Sevier, Washington, and Grand Wash fault zones, (Fig. 1; Pearthree, 1998; Stewart and Taylor, 1996; Zoback and Zoback, 1989). Movement along the Hurricane Fault has produced hundreds to thousands of meters of vertical displacement in the Late Cenozoic (Koons, 1945; Powell, 1875). Since Precambrian time, there has been some kind of geologic discontinuity near the present boundary of the eastern Great Basin and western Colorado Plateau (Wannamaker et al., 2001). The normal faults in northwestern Arizona and southwestern Utah are considered by Huntoon (1990) to be located along long-lived zones of weakness, which were active as normal faults during Precambrian time, and were later reactivated in the reverse sense during Laramide compression and uplift. These fault zones were activated again as normal faults during late Cenozoic extension (Spencer and Reynolds, 1989). This idea of repeated structural inversions accommodated along individual faults has been invoked to explain the rollover of hanging-wall rocks toward the large normal faults of the western Grand Canyon region (Hamblin, 1965). However, more recent studies of rift systems worldwide have demonstrated the ubiquity of complex structure, including hanging-wall folds, associated with normal faults (Janecke et al., 1998; Schlische, 1995). Exposures of the Hurricane Fault in the Grand Canyon do not reveal structure suggestive of an early history as a reverse fault overlain by a contractional monocline (Raucci, 2004). Thus, the reactivation scenario of Huntoon (1990), although often cited, may not be strictly applicable to the Hurricane Fault. In this case, the hanging-wall structure associated with the Hurricane Fault can be interpreted solely in terms of normal-faulting processes (Raucci, 2004). The modern Hurricane Fault separates the higher Uinkaret Plateau to the east of the fault from the lower Shivwits Plateau to the west. Fault movement, displacing Paleozoic to Quaternary rocks, has produced a steep, imposing curvilinear escarpment. Total displacement changes along strike with the largest displacements to the north and lowest offset to the south. The Hurricane Fault cuts into the Colorado Plateau near the Utah-Arizona

The Hurricane Fault has displaced strata ranging in age from Permian to Quaternary along the field trip route (Fig. 3). Paleozoic and Mesozoic strata have gentle northeast dips except near the fault, where strata typically dip moderately to the east in rollover anticlines, reverse-drag flexures, or monoclinal flexures depending on the terms one uses. The strata on the downthrown side dip toward the fault (up to 20° in some places, but generally 8°–12°; Billingsley and Workman, 2000). The escarpment relief along the Shivwits section varies from ~250 m at the Anderson Junction section boundary (Fig. 1) to ~200 m at the Moriah Knoll basalt flow. The relief is ~240 m at the escarpment convexity at the southern end near Twin Butte. Quaternary and Tertiary basalts, stream deposits, and colluvium cover much of the landscape along the fault. The basalts overlie the slightly tilted Paleozoic rocks forming an angular unconformity; and the basalts are faulted, though not as much as the underlying strata (Billingsley and Workman, 2000). Paleozoic marine carbonates, evaporates, and continental clastic rocks are exposed in the escarpment and as outcrops in Hurricane Valley. The following lithologic descriptions from Sorauf and Billingsley (1991) are listed from oldest to youngest (Fig. 3). The Toroweap Formation consists of three members: the Seligman Member, generally a slope former, is composed of pale red and reddish to yellowish-brown, fine-grained, calcareous sandstone, with some white to pink, granular gypsum and black, earthy dolomite. The Brady Canyon Member consists of tan, gray, and brown cherty, fossiliferous marine limestone. It is very resistant and forms cliffs and ledges. The Woods Ranch Member is composed of gray to black dolomite, tan sandstones, gypsum and gray to pale-red siltstone, which usually forms slopes. The overlying Kaibab Formation consists of two members: The Fossil Mountain Member, a light-gray, thick-bedded, cherty and sandy limestone that is a distinct cliff former at the top of the escarpment; and the Harrisburg Member, composed of light red to light-gray gypsum, limestone, and dolomite with minor amounts of red and gray siltstone and sandstone. The Harrisburg Member forms slopes with projecting carbonate ledges and is usually stripped back by erosion from the top of the escarpment. Where protected from erosion, remnants of Moenkopi Formation sandstone, siltstone, and shale red beds may be seen overlying the Harrisburg Member. The Mesozoic and Cenozoic sedimentary and marine units that overly the Moenkopi have been removed in much of the field trip area by erosion (Hintze, 1980; Billingsley and Workman, 2000; Billingsley and Wellmeyer, 2003). There are several late Tertiary and Quaternary basalt flows visible along the trip route that preserve some of the Triassic, Jurassic, and portions of the Cretaceous section. These are discussed in more detail in the road

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment

451

114° W

Cedar City

Cree k

N

Junction

ML 5.8 (1992) Utah Arizona

r

Cottonwood Canyon Fig. 6

Salt Lake City

Fredonia ISB

Shivwits

Main Street

Map area

Whitmore Ca nyon

Washington

Grand Wash

Mesquite

St.George

Mt. Trumbull

?

Flagstaff Phoenix

25 km Scale

122°W

45°N 105°W

45°N

Basin and Range

119°W 32°N

Figure 1. (A). Quaternary normal faults in northwestern Arizona and southwestern Utah, compilation adapted from Scarborough et al. (1986), Hecker (1993), and Pearthree and Bausch (1999). Significant recent earthquake epicenters (stars) and the sections (bold font) of the Hurricane Fault (Pearthree, 1998) are shown. Cottonwood Canyon, the site of recent seismic hazard assessment work on the Hurricane Fault in Arizona, is located north of the Shivwits-Anderson Junction boundary (Stenner et al., 1998). The Intermountain Seismic Belt (ISB) is a zone of earthquake activity extending through the Intermountain West from northwestern Montana south to Utah, southern Nevada and northern Arizona. The approximate boundaries of the ISB are shown in the inset.

Kanab

Toro wea p

37°N

Hurricane

Hurrican e

St. George

Anderson

Nevada

~M6 (1902)

Sevie

Ash

Zion National Park

Figure 2. Digital elevation model/hillshade map of the southwestern United States showing the Colorado Plateau and Basin and Range physiographic provinces. The dashed white box shows the extent of Figure 1.

°

Colorado Plateau

106°W 32°N

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log. These are seen capping some of the higher terrain and are evidence of topographic inversion. The stratigraphy exposed in the escarpment is similar along the Whitmore section south of the Shivwits section, but erosion by the Colorado River and tributary streams has exposed the lower Paleozoic section in and near lower Whitmore Canyon. Several middle Pleistocene basalt flows show vertical displacement of 23 m (Huntoon, 1977), middle to late Pleistocene basalt flows have estimated displacements of 6–20 m (Fenton et al., 2001; Stenner and Pearthree, 1999). FIELD TRIP ROAD LOG FOR THE SOUTHERN SHIVWITS SECTION OF THE HURRICANE FAULT ZONE, ARIZONA The Day 1 portion of the road log is 23 mi long starting just over the Utah State Line (Fig. 4A). The route continues onto the

Unit

Penn

Permian

Chinle Frm Moenkopi Frm

Triassic

Moenave Frm

Age

Arizona Strip, an isolated section of the State of Arizona not accessible from the south because there are no roads or bridges that cross the Grand Canyon and Colorado River between Marble Canyon and Hoover Dam, a distance of over 180 mi (290 km). There are good exposures of upper Paleozoic sedimentary rocks, capping late Tertiary basalt flows, fault scarps, and stream terraces and alluvial fans influenced by neotectonism along the route. There are no stops during this leg of the road log. Day 1 Road Log Cumulative mi (km) 0 0.2

(0) (0.3)

1.9

(3.1)

4.0

(2.4)

5.6

(9.0)

Thickness (m) Rock Type

Springdale SS Mbr

18-35

Whitemore Pt Mbr

0-20

Dinosaur Cyn Mbr

25-120

Petrified Forest Member

80120

Shinarump Member

0-50

upper red member

~150

Shnabkaib Member

~100

middle red member

~90

Virgin Limestone

~40

lower red member +/- Timpoweap Mbr

~170

7.7

(12.4)

Kaibab Limestone

75-260

8.1

(13.0)

+/- White Rim SS

0-60

Toroweap Formation

30-150

Hermit Shale

0-30

8.7

(14.0)

Queantoweap/ Esplanade / Coconino Sandstone

300-380

+/- Pakoon Fm.

0-90

Callville Limestone

60-275

Figure 3. Generalized stratigraphy for the Arizona portion of the Hurricane Fault region north of the Colorado River (Hintze, 1980, 1988; Lund et al., 2002). The relationship of late Tertiary and Quaternary basalt flows to the underlying stratigraphy is not shown.

Description Start at the I-15 Bloomington Exit #4. Drive east on Brigham Road. At 9 o’clock is an excellent exposure of the Shnabkaib and upper red members of the Moenkopi Formation. The unit capping the ridge is the Shinarump Conglomerate Member of the Chinle Formation. Turn right at intersection with River Road. Proceed south toward Arizona and the Moenkopi Terrace. At 12 o’clock, in the far distance, note gypsum-mining activity in the Harrisburg Member of the Kaibab Formation. Utah-Arizona state line. This area is partially covered by Quaternary alluvial deposits, including young deposits along active washes and eroded remnants of terraces and alluvial fans (Billingsley and Workman, 2000). At 9 o’clock, Pine Valley laccolith at the skyline. The high ridge to the southwest is Mokaac Mountain, a large ridge of Moenkopi Formation capped by late Tertiary basalt (Hamblin, 1970). Much of the slopes of Mokaac Mountain consist of large Quaternary landslides (Billingsley and Workman, 2000). A strand of the Washington fault zone, a major down-to-the-west normal fault, is at the base of steep, linear cliffs formed in resistant beds of the Kaibab and Toroweap Formations at 12 o’clock. The Washington fault zone extends from the town of Washington, Utah, southward for 60 km. The western strand, which is most apparent from this location, has been named the Mokaac Fault (Billingsley and Workman, 2000). Paleozoic rocks are displaced by ~200–400 m across the Mokaac Fault in this area, displacement increases to the northeast. Total displacement across the Washington fault zone to the east

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment

Figure 4 (continued on following two pages). Field trip route maps created from the U.S. Geological Survey St. George, Littlefield, and Mount Trumbull 1:100,000 topographic quadrangle maps.

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Figure 4 (continued).

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment

Figure 4 (continued).

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13.1

(21.0)

14.6

(23.5)

14.8

(23.7)

15.5

(24.9)

16.9

(27.2)

18.5 18.9

(29.8) (30.4)

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(34.1)

21.6

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is 100–600 m, also generally increasing to the north in this area. Quaternary activity of the Washington Fault has not been studied in detail, but late Quaternary alluvium and colluvium is displaced several meters at a number of locations (Billingsley, 1993d). The linear, steep escarpments along portions of the fault zone suggest substantial Quaternary activity (Menges and Pearthree, 1983). The road crosses the Mokaac Fault in this area although the fault zone is not exposed. Mokaac Fault to left, steep colluvial and alluvial deposits show some displacement. The road roughly parallels the main strand of the Washington fault zone for several miles. The Kaibab Formation crops out in the upper cliffs. The Harrisburg Member of the Kaibab Formation is eroded, whereas the Fossil Mountain Member is a cliff-forming unit. Red and yellow beds in the lower cliff are part of the Toroweap Formation (Billingsley, 1993d). We are crossing exposures of the Harrisburg Member and in some places red beds of the overlying Moenkopi Formation. Equivalent units on the upthrown side of the fault are high above us. Total vertical displacement across the Washington fault zone here is ~300 m (Billingsley, 1993d). Exposures of steeply dipping colluvial deposits in the road cuts indicate we are quite close to the fault zone but are still probably on the downthrown side of the fault. Seegmiller Mountain ahead; note white Shnabkaib member of the Moenkopi Formation. Junction with Seegmiller Mountain Road; take the right fork. Seegmiller Mountain is capped by a basalt flow over the Moenkopi Formation. Different ages have been reported for the flow, 2–3 Ma/ K-Ar, (Reynolds et al., 1986) 4–5 Ma/K-Ar (Wenrich et al., 1995), and 3.5–4.2 Ma/40Ar/ 39 Ar (Downing et al., 2001). Road roughly parallels the Washington fault zone on the upthrown side of the fault. We cross the Washington fault zone in this area. The ridgeline to the left of the road is capped with the Seegmiller Mountain basalt, which is displaced ~100 m here. Displacement of underlying Triassic strata is similar, implying that displacement on the Washington fault zone began less than 4 m.y. B.P. (Billingsley, 1993d, Downing et al., 2001). The Shivwits Plateau dominates the view to the south. The Shivwits Plateau is the hang-

22.9

(36.9)

ing-wall block of the Hurricane Fault and the footwall block of the Grand Wash Fault to the west. The Grand Wash Fault and its associated cliffs form the western margin of the Colorado Plateau at this latitude. The Shivwits Plateau is cut by a number of normal faults with much less displacement than either the Grand Wash or Hurricane Faults, including the Washington fault zone. First night’s camp on right (UTM 12S 271996 4072686). End of Day 1.

Day 2 This section of the road log covers the area of the Shivwits Plateau between the Washington and Hurricane Faults (Fig. 4B). The route crosses the Washington and Sunshine faults; the variations in amount and style of displacement can be observed. Stop 1, mile 40.1 (64.6 km) is an excellent overview of the Hurricane Fault zone and lower Hurricane Valley. Stop 2 is located near the relay ramp where the Moriah Knoll basalt crosses the Hurricane Fault, the lowest portion of the Hurricane Escarpment along the Shivwits section. Stop 3 is at the site of a paleoseismic trench excavated across the Hurricane Fault. Stop 4 is at the southern end of the Shivwits section where there are good exposures of the Holocene and Pleistocene alluvial fans and colluvial deposits on the Hurricane Escarpment. Road Log Cumulative mi (km) 22.9

(36.9)

23.8

(38.4)

24.2

(39.0)

Description Leave camp at 8:30 a.m., turn right on to road and proceed south. Bureau of Land Management sign identifying the Wolf Hole Valley. This small basin is on the downthrown side of the Washington Fault. The low escarpment in the distance at 10–11 o’clock is the continuation of the upthrown side of the Washington Fault. The high plateau to the west is Wolf Hole Mountain, which consists of Moenkopi Formation capped by late Tertiary basalt flows. Billingsley (1993d) obtained a K-Ar date of 3.1 ± 0.4 Ma from a basalt atop Wolf Hole Mountain. Near this location, Wolf Hole, Arizona, now defunct, consisted of a post office and general store to serve those living on the Arizona Strip, including the iconoclastic writer, Edward Abbey. Three small knobs, the Mustang Knolls, at 2–3 o’clock, are capped by basalts over the Moenkopi Formation.

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment 24.8

(39.9)

24.9

(40.1)

27.2

(43.8)

27.9

(44.9)

30.6

(49.6)

Continue straight at T intersection with Black Mountain Road. The upthrown block of the Washington Fault is evident at 9 o’clock, where there is ~60 m of total displacement (Billingsley and Workman, 2000). Bear left at Y intersection, road crosses the Washington Fault. High point in the road at Wolf Hole Pass with views to the east and south. The high escarpment in the distance (12–2 o’clock) is the Hurricane Cliffs. The Uinkaret Plateau, on the upthrown side of the Hurricane Fault, has numerous eruptive centers. The high, broad mountain in the far distance is Mount Trumbull, the remnant of an early Quaternary volcano. Left of Mount Trumbull is Antelope Knoll, another Quaternary volcano. Diamond Butte at 1 o’clock is composed of Moenkopi Formation capped by basalt dated at 4.3 ± 0.6 Ma/K-Ar (Billingsley, 1993c). Extensive Pliocene basalts capping buttes and mountains on or around the margin of the Shivwits Plateau imply that the land surface was generally formed on the Moenkopi Formation in the early Pliocene. In most places, the Moenkopi deposits have been completely removed and the surface of the Shivwits Plateau is now formed on the underlying Kaibab Formation. Erosion of the Moenkopi Formation must have provided abundant sediment to the Colorado River system in the past few million years. The road here is very close to the contact between the Permian Kaibab Formation to the north and the Triassic Moenkopi Formation to the south. In this area, the Moenkopi Formation filled a Triassic paleo-valley carved into the Kaibab Formation (Billingsley, 1991). Bear right at Y intersection as you enter Main Street Valley. Main Street Valley is bounded on the east by an obvious escarpment associated with the principal strand of the Main Street Fault. Total displacement across this eastern fault strand ranges from 50 to 120 m. Quaternary deposits and basalts are displaced in a few locations, but no detailed studies have been completed on this fault zone. Portions of the west side of the valley are also fault-bounded, and this part of the fault system has been labeled the Main Street graben (Billingsley, 1993a; Menges and Pearthree, 1983).

31.9

(51.3)

33.1

(53.3)

35.6

(57.3)

36.1

(58.4)

36.8

(59.3)

37.2

(59.9)

37.7

(60.7)

39.3

(63.3)

457

Turn left onto the Navajo Trail and proceed to the east across the principal Main Street Fault strand. We are now driving on the Harrisburg Member of the Kaibab Formation. The knoll in the far distance at 12 o’clock is Antelope Knoll (Fig. 5). The Fossil Mountain Member of the Kaibab Formation crops out in the lower parts of this valley. This is a prominent cliff-forming unit that you will see high up on the footwall of the Hurricane Fault. Hurricane Cliffs are visible in the distance at 12 o’clock. We cross the Sunshine Fault as we emerge into a valley. The Sunshine Fault has generated an escarpment at the southwestern end of this valley; total displacement is ~110–130 m, down-to-the-east (Billingsley, 1993a). The Sunshine Fault is likely an east-dipping normal fault, antithetic to the Hurricane Fault. It has the most linear and impressive escarpment of the secondary faults associated with the Hurricane Fault. The road is on a young alluvial fan that is not displaced by the Sunshine Fault. In the valley to the east there are several low scarps that are probably related to the west-dipping normal faults. Part of the Hurricane Cliffs Escarpment is visible from 11–3 o’clock. The Hurricane Fault is at the base of the escarpment. With mid-morning light, you may be able to see some low fault scarps formed in the colluvium/alluvium at the base of the cliffs that record late Quaternary displacement on the fault. The Pine Valley Mountains north of St. George, Utah, can be seen at 9–10 o’clock. A structurally complex portion of the Shivwits section (the Grandstand; Fig. 4) of the Hurricane Fault and the Navajo Trail crossing of the escarpment are seen at 11–1 o’clock. The modest escarpment we are crossing is associated with another minor, west-dipping fault strand. The road is built on rocks of the Moenkopi Formation that filled a paleo-valley in the Kaibab Formation. The Kaibab Formation crops out at the crest of the ridge. The ridgeline in the middle distance at 9–10 o’clock is capped with late Tertiary basalt dipping moderately toward the Hurricane Fault.

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Figure 5. Panoramic view at Stop 1 of the Hurricane Escarpment that covers much of the Shivwits section of the Hurricane Fault. Geographic locations include Black Rock Canyon just south of Cottonwood Canyon (A), a stepover on the Hurricane Fault, where the Navajo Trail (built on a relay structure, Peacock and Sanderson, 1991) crosses the escarpment, the Grandstand (C), Antelope Knoll (under the D), Moriah Knoll basalt cascade over the escarpment (E), and Moriah Knoll (under the F).

39.8

(64.0)

40

(64.4)

40.1

(64.6)

40.1

(64.6)

Large stock tank and causeway at the crossing of Hurricane wash. Crossing a low escarpment associated with a west-dipping fault. Basalt-capped Diamond Butte is visible at 2 o’clock. The axis of the rollover Hurricane monocline is somewhere in this vicinity (Billingsley, 1993b). Park on the south side of the road. Walk ~100 m south of the road to the crest of a low hill. This hill is capped with a thin layer of alluvium that was probably part of an alluvial fan that associated with the relict course of Hurricane Wash (Billingsley, 1993b). It is likely that some combination of displacement along the fault zone immediately to our west and associated drainage capture altered the course of Hurricane Wash to its present position. Stop 1 (UTM 12S 0291529 4066806, Grandstand 7.5′ quadrangle, T38N, R10W, NW/4 Section 1).

Stop 1—Overview of the Shivwits Section of the Hurricane Fault From this vantage point we can see much of the Shivwits segment of the Hurricane Fault. Down-to-the-west displacement across the Hurricane Fault has resulted in the formation of the Hurricane Cliffs, the prominent escarpment that dominates our view to the east (Fig. 5). The Hurricane Cliffs separate the higher Uinkaret Plateau to the east from the Shivwits Plateau on which we are standing. Two Quaternary eruptive centers on the footwall of the Hurricane Fault (Fig. 6) can be seen in the distance above the Hurricane Cliffs; the obvious volcanic cone is Antelope Knoll and less obvious is Moriah Knoll to the south (Billingsley, 1994a, 1994b). Stop 1 is located on the monocline in the hanging wall where strata on the downthrown side dip toward the fault. This increase in dip toward the fault has been suggested to be the result of reverse drag flexure due to decreasing fault dip at depth (Billingsley and Workman, 2000; Hamblin, 1965). The Hurricane Cliffs are capped by the Harrisburg Member of the Kaibab Formation (Fig. 3), but in many places, it has been stripped from the escarpment. The highest, steep cliff is formed

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment

Neotectonics of the Hurricane Fault The Hurricane Fault provides excellent exposures of displaced Quaternary alluvium and basalt flows for the evaluation of seismic hazard and discerning its neotectonic history. Lund and Everitt (1999) and Stenner and Pearthree (1999) all have identified displaced basalt and alluvium that indicate that the Hurricane Fault has been active throughout the Quaternary. Paleoseismic investigations of the Anderson Junction section (Fig. 1) discovered evidence of several Pleistocene and Holocene surface-rupturing earthquakes. Stenner and others identified a latest Pleistocene to early Holocene surface-rupturing most-recent event (MRE) at Cottonwood Canyon on the Anderson Junction section (just south of the Utah border, Fig. 1) with 0.6 m of vertical displacement (Stenner et al., 1998). Further trenching investigations at Rock Canyon, 4 km north of Cottonwood Canyon, revealed that the last three events had variable amounts of slip per event (Stenner et al., 2003). The MRE had an estimated 0.3–0.4 m net vertical slip, whereas the penultimate and pre-penultimate events together had ~2.7–3.7 m of vertical slip. Possible scenarios to explain the lower MRE offset at Cottonwood and Rock Canyons include the rupture of the Shivwits that propagated north into the adjacent southern Anderson Junction sections, or a separate rupture in the boundary between the two sections. The size of older fault scarps at Cottonwood and Rock canyons, along with estimates of earthquake recurrence intervals (5–100 ka) in the Basin and Range province (Stenner et al., 1998), suggest that larger slip-per-event (more than 0.6 m) is typical along this part of the Hurricane Fault. A paleoseismic investigation here along the Shivwits section of the Hurricane Fault revealed evidence of surface-rupturing late

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in the Fossil Mountain Member of the Kaibab Formation. Lower slopes on the cliffs consist of the Woods Ranch and Seligman Members of the Toroweap Formation, which bracket the cliff-forming Brady Canyon Member. Coarse, very poorly sorted colluvium covers much of the slope-forming units, especially the Seligman Member. Because of the structural complexity of the Hurricane Fault zone immediately across the valley from this vantage point, you may see all or parts of this sequence repeated several times. Most of the Shivwits segment is a large structural embayment between two prominent convex fault bends. On the hanging wall at the northern end of the Shivwits segment, you can see a prominent east-sloping butte (mentioned at road-log mile 39.3), where beds of the Triassic Moenkopi Formation are capped with late Tertiary basalt and all are tilted toward the Hurricane Fault. This butte is at a major convex bend in the trace of the Hurricane Fault similar to the State Line geometric bend (Stewart and Taylor, 1996). Immediately northeast of our overlook, the gravel road of the Navajo Trail can be traced up to the Hurricane Escarpment, where it ascends the surface of a ruptured relay ramp between overlapping strands of the Hurricane Fault. The Grandstand, seen south of the Navajo Trail, is a zone of multiple fault strands associated with a left stepover. To the southeast, we can see the Moriah Knoll basalt where it flowed across the escarpment (Fig. 6). This is discussed at Stop 2.

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Figure 6. Mosaic of NASA high-altitude aerial photography of the central part of the Shivwits section of the Hurricane Fault showing fault traces, slip-rate estimates, and field trip stops. The fault strands are from Billingsley (1994a and 1994b) and field reconnaissance. Faults are dashed where approximate or inferred, dotted where concealed. The relay ramp, south of the basalt flow, is evidence that fault linkage had occurred on this portion of the fault (Peacock and Sanderson, 1994). Shown is a compilation of the vertical surface displacement observations (the format is profile #/surface offset in m/slip-rate range in mm/ yr). Profile #7 is the Boulder Fan (outlined) trench site. The basalt flow displaced by the Hurricane Fault originated from Moriah Knoll.

Quaternary events (Amoroso et al., 2004). Mapping did not show any evidence of surface rupture of Holocene deposits; the only convincing evidence of tectonic displacement was found in late to middle(?) Pleistocene alluvial fans. Results using displacement of the Moriah Knoll basalt (Fig. 7), topographic profiling, surface dating, morphologic modeling of fault scarps, and observations

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Figure 7. (A) Geologic map showing the relation of the faults to the Moriah Knoll basalt (Qmb) and the mapped flow directions (heavy black arrows). Mapped surficial units: Pkh, Harrisburg Member of the Kaibab Formation; Pt-k, Permian Toroweap and Kaibab Formations undifferentiated; Qm1–3, from mid to late Pleistocene alluvium; Qy1, Holocene alluvium; Qco, Quaternary colluvium. The basalt flowed through a paleo-canyon, crossed the fault, and covered the relay ramp surface between the escarpment and the Pkf western ridge of the relay ramp until the ridge was overtopped and the basalt flowed further northwest (B). The basalt flow directions, estimated from flow thickness, are shown by heavy black arrows. A–A′ is the location of the cross-section C. The letters on the map (A through F, EF, WF) refer to locations discussed in the text. (B) Photograph looking NE toward the Hurricane Cliffs, note the relay ramp and basalt flow that crossed the escarpment, flowed across the relay ramp, to the valley floor. (C) Cross section A–A′ showing the estimation of maximum vertical displacement of the Moriah Knoll basalt.

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment in the paleoseismic trench all yield a slip-rate estimate of ~0.05–0.3 mm/yr during most of the Quaternary. Multiple time scale determination of the offset rate provides important constraints for our understanding of the rupture geometry, timing, displacement per earthquake, and earthquake hazard along the Hurricane Fault. The 2.3–3.4 m of surface displacement due to the MRE, ca. 10 ka, suggest that the Shivwits section is capable of producing M7 earthquakes with an estimated average recurrence interval of 9–15 k.y. These results will be discussed further at Stops 2 and 3. The neotectonic implication of the results in Amoroso et al. (2004) is that the western edge of the Colorado Plateau has been deforming at a roughly constant rate for at least the past million years and thus that Basin-and-Range style crustal extension is active along the plateau margin.

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Description Rocks of the Moenkopi Formation crop out on the left side of the road here. These sediments filled a NE-trending Triassic paleo-valley (Billingsley, 1993b).

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Turn right onto small dirt road. We are climbing onto older Quaternary fluvial deposits (Billingsley, 1993b) laid down by a northward-flowing stream in this valley. The hummocky part of the Hurricane Cliffs at 11–12 o’clock is where the Moriah Knoll basalt flowed across the Hurricane Fault zone (Figs. 7 and 8). The Temple Trail traverses the Hurricane Cliffs on the Moriah Knoll basalt. Mormon pioneers constructed this trail to bring large Ponderosa Pine logs from Mount Trumbull to St. George, where they were used as roof beams in the construction of the Church of Latter Day Saints temple. The trail up the basalt flow is not passable by motor vehicles. After descending the low ridge to the south, we are driving on low terraces and the flood plain of a sizable, unnamed north-flowing tributary of Hurricane Wash. Gate. Another gate; turn right at Y intersection. The Boulder Fan trench site is visible in the middle distance at 12 o’clock (Stop 3 on Figs. 4B and 6). The large rockfall boulder on the fan is responsible for its name. Twin Butte

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Figure 8. Variation in escarpment height along the Shivwits section; the x-axis is the distance along the escarpment, the local strike of the escarpment/Hurricane Fault is shown. The northern end of the Shivwits–Anderson Junction section boundary is at km 5, approximately where the fault strike changes 55°. The change in escarpment height shows the variation in total slip along the section. The basalt locale is a geometric boundary and may be a segment boundary where total slip decreases and stress is distributed over several fault strands in the area (see Zhang et al., 1991).

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is visible on the hanging wall of the Hurricane Fault at 2 o’clock. A drilling rig is visible in the middle distance at 2 o’clock. This is called the Dutchman Prospect, and it is being drilled into a surface anticline between the Hurricane and Sunshine fault zones. Potential source rocks are Paleozoic and Precambrian shale. Drilling began in 1998 using an old-fashioned cable tool rig. Because of hard drilling and lack of progress, the rig was converted to rotary drilling. As of mid-2001, the total depth was ~3200 feet. The drilling rig was still there as of May 2005. Turn left at the gate in the fence and head toward the relay ramp/basalt flow. Crossing Holocene debris flow, note the boulder-rich character and the distance this deposit extends into the basin floor. Stop 2 (UTM 12S 298411 4060533, The Grandstand 7.5′quadrangle, T38N, R9W, nw/ 4 section 27) at a metal stock tank.

Stop 2—Moriah Knoll Basalt The Moriah Knoll is one of several late Quaternary eruptive centers that are part of the Uinkaret Volcanic Field (Billingsley, 1994a, 1994b). The Moriah Knoll volcano erupted basalt radially onto the footwall of the Hurricane Fault (Fig. 6). From mapping of flow thickness and texture (Fig. 7A), it appears that the basalt flowed south along a developing relay ramp, accumulated against the relay ramp ridge, overtopped the ramp to the west (Fig. 7B) and flowed down onto the hanging wall (Amoroso et al., 2004). Displacement along several fault strands has subsequently offset the flow. The various exposed remnants of the Moriah Knoll basalt flow, on both sides of the Hurricane Fault, were sampled and correlated using X-ray fluorescence spectrometry geochemical analysis of trace elements (Amoroso et al., 2004). A sample collected from the flow in a paleo-canyon on the footwall gave a 40 Ar/39Ar date of 0.85 ± 0.06 Ma. Billingsley (1994a) measured ~130 m displacement of the Moriah Knoll basalt flow by the Hurricane Fault by estimating total topographic offset along the fault strands. Using Billingsley’s displacement estimate yields a 0.15 ± 0.02 mm/yr slip rate. Amoroso et al. (2004) also estimated the total topographic displacement from where the flow crossed the paleo-escarpment to the base of the flow farthest from the fault (Fig. 7E–7F). We considered the likely paleo-topographic elevation changes across the fault strands. Cross section A–A′ (Fig. 7A) is oriented approximately along the presumed basalt flow axis. The displacement is based on the average elevation of the base of the westernmost exposed part of the flow. The difference in altitude from the base of the basalt in the paleo-canyon to the exposed base of the flow at its westernmost edge is 204 m. This is probably a minimum displacement, and yields a slip

rate of 0.24 ± 0.02 mm/yr. If the 2° slope in the paleo-canyon is extrapolated 1.4 km to the westernmost basalt flow, ~50 m of this offset estimate may be the original topography. The remaining 154 m of displacement yields a slip rate of 0.18 mm/yr. From these estimates, we infer a long-term slip rate of 0.15–0.24 mm/ yr for this part of the Shivwits section. The reported range is due to the uncertainty in the basalt age, inferences about the paleotopography, and error in measuring the displacement using the topographic map (± one contour, 20 ft or 6.1 m). There could be unaccounted for displacement along fault strand WF (Fig. 7) or on faults along the westernmost part of the flow (E). Alternatively, the westernmost part of the flow could be buried by alluvium, been eroded away, or never emplaced. Fault Segmentation The idea of fault segmentation is based on observations that often only part of a long fault zone ruptures during a large earthquake (Crone and Haller, 1991; Schwartz, 1988; Schwartz and Coppersmith, 1984). Geometric discontinuities such as abrupt changes in fault strike, gaps, and stepovers, zones of increased structural complexity, and behavioral discontinuities including change in slip rate or displacement sense may act as rupture barriers (dePolo et al., 1991). The Hurricane Fault is more than 250 km long (170 km in Arizona) and has changes in strike, stepovers, and complex structure; therefore, it probably ruptures in segments (Stewart and Taylor, 1996). The major Hurricane Fault section boundaries were originally defined by Pearthree (1998) using large changes in fault geometry and displacement across the fault (Fig. 1). The boundaries of the Shivwits section are defined by a major convex fault bend 10 km south of the Utah border (north boundary) and a zone of structural complexity and diminished displacement just west of Mount Trumbull (south boundary). The next fault section south is the Whitmore Canyon section, which extends south to the Colorado River; the Anderson Junction section lies north of the Shivwits section and continues north into Utah. Evidence of Fault Segmentation along the Shivwits Section A change in strike of the Hurricane Fault at the Moriah Knoll basalt flow (Fig. 8) along with substantial fault complexity suggests that this area may be a geometric boundary (Zhang et al., 1991). The Moriah breached relay ramp area was probably a segment boundary prior to linkage. Relay ramps accommodate displacement between normal fault zone segments; breached relay ramps are produced when the fault segments join to form a linked fault (Peacock and Sanderson, 1994). Evidence of the fault linkage is shown in Figure 7A. The western fault (WF) of the relay ramp in Figure 7A (striking SW) has increased displacement to the south, while the eastern fault (EF) north of the canyon has increased displacement to the north (Billingsley, 1994a). The EF appears to have breached the relay ramp, which has consequently tilted toward the east. Additionally, this portion of the fault shows less total displacement than more linear parts of the fault (Fig. 8). Peacock

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment and Sanderson (1991) noted a decrease in total fault displacement near relay structures, accommodated by bending and/or tilting of the relay structure or folding of wallrock. The Shivwits Escarpment is presumed to be a time-integrated displacement record of hundreds of earthquakes; lower total displacement at this location suggests this area is a fault segment boundary. A second example of lower fault displacement along the Shivwits section is the Twin Butte area (Figs. 6 and 8) and adjacent Diamond Butte. These landforms lie near a convexity on the Hurricane Fault. These features may be a rupture barrier, and Janecke (1993) suggested this type of feature may be evidence of a segment boundary. Changes in fault strike and escarpment height suggest that the Moriah Knoll basalt locale is another geometric boundary and might be a segment boundary. The fault strike changes from N20°E north of Twin Butte, to N15°W at the basalt locale, to N35°W south of the State Line geometric bend (Fig. 8). The escarpment is 470 m high at the State Line geometric bend. Farther south along the escarpment (Fig. 6), we can see another salient of the Hurricane Fault at the basalt-capped mesas on the hanging wall, Twin Butte, and to the west, Diamond Butte. These buttes preserve remnants of the easily eroded Moenkopi Formation on the hanging wall because of the resistant cap of late Tertiary basalt. Near Twin Butte, the escarpment is 450 m high. The escarpment height is lowest (250 m) where the Moriah Knoll basalt crossed the fault zone (Fig. 8). The overstepping faults depicted in Figure 7A are part of a transfer zone (Peacock and Sanderson, 1994). The eastern fault shows increasing displacement to the north, and the fault to the west has increased displacement to the south (Billingsley, 1994a). We suggest that the minimum in escarpment height at the Moriah Knoll basalt is related to fault geometry. The basalt flowed down the relay ramp between these faults where total displacement decreases and stress is distributed over several fault strands in the area (Zhang et al., 1991). The next stop, the Boulder Fan trench site, is another part of the fault where hard linking of fault strands along the fault zone has isolated a smaller fault strand. Road Log Cumulative mi (km) 48.9

(78.7)

Description Fault scarps along the base of the Hurricane Cliffs are apparent at 11 o’clock, especially when the sun is low in the sky. Alluvial fault scarps that record the past few faulting events are very common along the base of the cliffs from this area south to Twin Butte. These scarps are formed in steep late Quaternary alluvial fans or colluvial slope deposits composed of rather coarse, poorly sorted gravels. Morphologic analyses based on diffusion modeling suggest an early Holocene

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to late Pleistocene age for these fault scarps (Amoroso, 2001). The south-dipping Moriah Knoll basalt surface in the relay ramp is at 9 o’clock. The basalt obviously flowed south along the relay ramp, but it also has likely been tilted southward because of increasing fault displacement to the south subsequent to its eruption. Boulder Fan and trench are in view at 11 o’clock. The drainage that cut into the cliffs south of the Boulder Fan is the location of a fault splay that continues south then southwest behind the escarpment. Proceed through gate at Merchant Tank and bear left. The U.S. Fish and Wildlife Service released a number of juvenile condors into the wild atop this section of the Hurricane Cliffs in 1999 as part of efforts to ensure the survival of this largest North American bird species. The birds remained in the area through most of 2000, but eventually flew off to join their colleagues in the eastern Grand Canyon area. It is rumored that they left because either the geology is more interesting or the pickings are better in Grand Canyon. Turn left onto a dirt track, follow it toward the Sheep Pockets Reservoir (a large stock tank), proceed around the tank to the left and follow the track up an increasingly steep alluvial fan to the base of the cliffs. Park near the large boulder. Stop 3—Boulder Fan trench (UTM 12S 0298576 4054388, Russell Spring, 7.5′ quadrangle, T37N, R9W, NW/4 Section 15); 75 min including lunch.

Stop 3—Boulder Fan Trench This stop is near the apex of the Boulder Fan where we excavated a 70-m-long trench (Fig. 9) across the Hurricane Fault at the base of the Hurricane Escarpment at profile line 7 (Fig. 6) in 2000. The southern wall of the trench is remarkably intact illustrating the partly consolidated nature of the debris flow sediments. Topographic profiles across the fault scarp show 4–4.6 m of far-field vertical surface displacement of the fan surface. The age of the alluvial fan surface was estimated to be late Pleistocene using a calibrated carbonate rind proxy (15–75 ka; Amoroso et al., 2004). Carbonate rinds, an accumulation of pedogenic carbonate on clasts, have been previously used for cross-correlation of landforms as well as a quantitative indicator of soil age. Rind thickness of clasts varies within a soil profile (usually 1 m deep). Samples are collected at 10 cm intervals and a maximum thickness value is determined for that profile. Surface ages can then be estimated using a rind thickness to surface-age proxy (Amoroso, 2005).

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Figure 9. (A) Topographic map of the Boulder Fan produced using total station surveying. Elevations and locations measured relative to an arbitrary datum, 1 m contour interval. Arrow indicates the 6 m boulder in the photograph. (B) The Boulder Fan; the large boulder (arrow) in the foreground likely fell from the cliff-forming Fossil Mountain Member of the Kaibab Formation in the upper part of the escarpment. Beneath the Fossil Mountain Member are members of the Toroweap Formation. The slope-forming Woods Ranch Member and the cliff-forming Brady Canyon Member and the slope-forming Seligman Member make up the lower three-fourths of the escarpment. (C) The rose diagram shows results of 161 measurements of clast imbrication (long-axis orientation) of Unit 40. Clast-source direction vector was 287° ± 10° (95% confidence interval); the results suggest a bimodal distribution.

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment Paleoseismic Trench Investigation This trench revealed complex stratigraphy in the hanging wall associated with fault deformation, rotation of strata adjacent to the fault, and erosion of the fault scarp. We believe there is solid evidence for two faulting events since deposition of the fan sediment; the youngest event occurred ca. 10 ka. Each faulting event involved at least 1.5 m of surface displacement, suggesting that they were associated with large earthquakes that ruptured much or the entire Shivwits segment of the Hurricane Fault. The following is a summary of major features of the fault zone. See Amoroso et al. (2004) for a more extensive discussion and description of stratigraphic units. The trench site is on a large, late Pleistocene alluvial fan (the Boulder Fan; Figs. 9A and 9B). The fan shows vertical surface displacement of ~5 m. The fan surface shows only minor surface erosion by small drainages, but the fan is bounded to the north and south by larger drainages that are incised 10–15 m (Fig. 9A). The trench site has several 1– 6 m diameter boulders on the surface that were almost certainly emplaced by rockfall from the Hurricane Cliffs to the east (G.H. Billingsley, 2000, personal commun., observed similar rockfalls during the 1992 earthquake). General Trench Stratigraphy The stratigraphic units are numbered from oldest to youngest and grouped by their association with the most-recent event (MRE) and penultimate (PE) earthquakes. The stratigraphy of the footwall consists of a series of ~0.5- to 3-m-thick debris-flow deposits (Fig. 10). These deposits accumulated across the fan surface until normal faulting ruptured them. The ruptured alluvial fan surface created scarps that were subsequently eroded, creating colluvial deposits on the hanging wall. During the trench investigation, we sought to correlate stratigraphic units across the fault. We examined debris flow deposits on the western edge of the trench to test the correlation of units across the fault. The implication of correlating these units is that there has been ~4.6 m of vertical surface offset of the alluvial fan surface across the fault zone. Fault-Zone Stratigraphy Descriptions of the units adjacent to the fault and those genetically related to the fault zone are followed by interpretation of their style of deposition. A summary of the correlation of units across the fault and estimation of amounts of unit displacement is given here; see Amoroso et al. (2004) for more detail. The debris flow deposits are the dominant facies in the fan, they consisted of sandy to silty-sand gravels and cobbles, primarily matrix supported. The buried debris flow deposits showed significant soil development at the upper portion of the deposit. The uppermost silty zones in Unit 3 and 4 may be eolian additions to the surface of the debris flows. Stratigraphic units exposed in the hanging wall consist of clast-supported fluvial gravel (Unit 40) that may have filled a trough along the base of the scarp, and a mantle of toeslope colluvium derived from the scarp.

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Unit 13 on the upthrown side appears to be correlative with Unit 7 on the downthrown side when comparing consistence, texture, and structure. Unit 13 has more carbonate accumulation than Unit 7; this may be related to variation in soil development down slope. Tracing the top of Unit 7 eastward toward the fault zone, the slope on the top of the unit decreases east of the 18 m mark (Fig. 10) and becomes essentially horizontal between 34 and 39 m. From this relation, we infer some amount of back-tilting of Unit 7 that would result in the observed scarp height being greater than the measured surface displacement. The trough that formed along the base of the scarp is mostly filled by sediment derived from erosion of the scarp and possibly from adjacent drainages, however. Structural relations of the units to the fault zone are shown in Figure 11. The easternmost fault is marked by weak shearing of the fabric in Unit 1 in fault contact with Unit 30 and by several rotated and broken clasts. Two sedimentary packages, Units 40, 41, 42, 45, 46, and 47 (taken together) and Unit 21 (Fig. 11), have characteristics that suggest they are colluvial wedges from postrupture erosion of the fault scarp and subsequent deposition on the downthrown side. The western fault is marked by a difference in color and consistence between Units 9 and 21 compared with Units 30, 31, 32 and 34; small fissures, faint shear fabric, and rotated clasts were observed along the fault trace. Unit 9 is likely a debris-flow deposit. It appears correlative with a texturally similar debris-flow deposit Unit 7 (Fig. 10) and probably represents the pre-faulting fan surface. The clast fabric dip is somewhat less than in Unit 21, which likely represents discordance between the dip of the original fan sediments and the overlying fault related colluvium. While it is possible that Unit 9 is a colluvial wedge, the observations by Ostenaa (1984) that colluvial wedge thickness was approximately half of the free-face height makes this unlikely. Unit 40 is a framework gravel and cobble deposit interpreted to be a fluvial deposit because of clast imbrication and the paucity of matrix. The clast imbrication vector of 287° ± 10° indicates the clasts were transported over the scarp perhaps from a paleo-channel near the present southern channel (Fig. 9). The detailed map of the Boulder Fan (Fig. 9A) shows that for fluvial transport of the clasts from the southern channel along the fault scarp the stream would have to bend ~120° and flow up slope. Fluvial input from the northern channel along the fault scarp is not likely, based on the imbrication data. Based on the thickness and position of the colluvial wedges, the eastern fault appears to have accommodated most of the movement during the MRE. The western fault is clearly defined by shear fabric and accumulation of additional soil carbonate low in the trench. The zone of cemented and sheared gravels between the faults is deformed, and a depression apparently formed at the top of this zone during both the PE and MRE. There was an estimated 15–25 cm of movement on the western fault in the MRE, based on the displacement of Unit 35 compared to Unit 21. Interpretation of the fault zone stratigraphy based on these logs indicates that there were at least two, and less likely three,

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Figure 10. Log of the south wall of the Boulder Fan trench. The alluvial fan on the upthrown side is a series of debris flows with a thin veneer of slope wash materials (Unit 51). The downthrown side shows units derived from the degrading fault scarp and a fluvial unit that encroached on the fan soon after the most-recent event (MRE). Considering texture, consistence, and structure, Unit 13 on the upthrown side appears correlative with Unit 7 on the downthrown side, which appears to be back-tilted within ~10 m of the fault zone. The buried soil zone in Unit 7 (between 34 and 36 m marks), the clast fabric of the eastern part of Unit 7 (between 30 and 36 m marks), and the organic-rich layer show that these surfaces have tilted back toward the fault. We did not see antithetic faulting west of the fault zone. The eastern fault showed greater movement during the MRE than the penultimate event. There was ~15–25 cm of MRE movement along the western fault. The zone of cemented and sheared gravels between the faults may be a small graben that formed during the penultimate event. Projecting the tops of Units 13 and 7 to the fault zone suggests there has been ~4.6 m of vertical surface displacement. Optically stimulated luminescence samples discussed in the text were collected from Units 3 and 4 near the top of the trench (67.5 m mark).

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18

Elevation, m

20

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20A 42 3

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9 Weak shear fabric Gradational contact - enhanced carbonate cementing from here to fault B

Burrow

Figure 11. Detail of the fault zone. We deepened the trench at the fault zone to further investigate the nature of the faulting and expose more of Unit 9. Unit 9 may be a debris-flow unit that may show the effect of fault drag. Alternatively, Unit 9 may be evidence of a third event. Support for the latter model comes from the steepened clast fabric (dip shown in brackets) near the fault, which is similar to the PE wedge (Unit 21) and has an increase in clast content away from the fault. Ostenaa (1984) suggested that wedge thickness is approximately one-half of the free-face height. We did not find the lower contact of Unit 9, there was a subtle change in texture near the base of the trench. If Unit 9 is a colluvial wedge, it would be a 2+ m thick wedge, which suggests 4+ m single-event surface offset. The location of the radiocarbon samples from Units 21 and 33 are outlined.

[12°]

[47°]

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[35°]

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[35°]

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Base of trench

events that ruptured the Boulder Fan. A total displacement of the fan surface of ~4.6 m suggests average displacements of 2.3 m per event in two events or ~1.5 m per event for three events. These are minima because we measured 5.5 m of throw at the fault zone. The two-event model is preferred because the evidence for a third colluvial wedge is equivocal. Bulk samples were collected from two locations in the fault zone to constrain the timing of earthquake events; the fissure sample provided a useful amount of dateable material (Fig. 11). Sampling the fissure would likely estimate the minimum limiting age of the MRE rupture (McCalpin and Nelson, 1996). The result was a radiocarbon age of 9910 ± 210 yr B.P. for the MRE. The corrected 2σ calendar age (using Calib 4.3, http://depts.washington.edu/qil/ calib) of the sample is 9300 +1070/−430 cal yr B.P. The resulting age is considered a minimum age of the MRE, assuming that the fissure filled soon after surface rupture. Geomorphology of the Hurricane Fault Area This stop is an excellent location to overview the general geomorphology of the Hurricane Escarpment. The trend of the Hurricane Fault, and therefore the escarpment, is generally north-

0

1m

south with portions oriented northwest or northeast (Figs. 1 and 6). The fault trend is sinuous and results in salients (convex to the west) and re-entrants (concave to the west) that have been described as geometric segment boundaries of the Hurricane Fault by Stewart and Taylor (1996). Stratigraphy plays an important role in the escarpment morphology. Because of the lower resistance to erosion of the Woods Ranch and Seligman members along the base of the cliffs, there has been significant cliff retreat along the escarpment as indicated by the fault trace, located tens to a few hundred meters west of the main escarpment. Where the Woods Ranch Member (gypsum, gypsiferous siltstone and silty sandstone) outcrops in the escarpment appears to play an important part in controlling the distance from the fault trace to the escarpment. Stenner and Pearthree (1999) noted that where the softer Woods Ranch Member comprises a significant portion of the escarpment, cliff retreat appears greater and the fault trace is farther from the escarpment. Varnished, older colluvial deposits are evident along the escarpment in areas that generally lie away from the active landslide chutes and drainages. In the Cottonwood Canyon area

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(Fig. 1) there are several relict colluvial boulder deposits that are offset substantially by Quaternary faulting (Fig. 12) suggesting they are of middle to late Pleistocene age. Studies of older desert hillslope deposits have interpreted these deposits as indicators of increased colluvium production during late Pleistocene pluvial periods (Gerson, 1982; Whitney and Harrington, 1993), though some work suggests they may be much older (Friend et al., 2000). The alluvial fans along the base of the escarpment were likely created by debris flow action stripping the boulder deposits during wetter-to-dryer climate transitions (Bull, 1991; Pearthree and Bausch, 1999; Pederson et al., 2000). We can make indirect estimates of the age of these deposits in the Cottonwood Canyon area (Fig. 12). Scarps visible in the lower part of the escarpment are ~20 m high. Using conservative estimates of 0.1–0.3 mm/yr for fault slip rate made for the Anderson Junction and Shivwits sections of the Hurricane Fault (Stenner and Pearthree, 1999; Amoroso et al., 2004) yields an age of ca. 70–200 ka for the faulted colluvial slopes. Alluvial fans are common along the Hurricane Escarpment and extend well into the Hurricane Valley. Mapping identified five distinct ages of alluvial fans from middle (?) to late Pleistocene (Qm1) to late Holocene (Qy2) (Amoroso, 2001). Age estimates were made using soil development including a locally calibrated carbonate-rind thickness proxy (Amoroso, 2001). The method is further developed for use more widely in the southwestern deserts of the United States (Amoroso, 2005). The Pleistocene (Qm1 to Qm3) alluvial fans are primarily composed of stacked debris flow deposits with minor slopewash and stream alluvium. These fans are composed of partly to well-consolidated matrixsupported deposits of gravels, cobbles, and boulders, and are

confined to locations near the base of the escarpment. Pedogenic carbonate accumulation is common, from discontinuous carbonate rinds (Stage I) in Qm3 fans to Stage III soil carbonate in Qm1 fans. Buried soils formed in the debris flow materials were observed in the trench on the Boulder Fan (Fig. 10), suggesting considerable time could elapse between subsequent debris flow deposits. The modern surface alluvium contains a significant silt fraction coming from the frequent dust-laden windstorms. The buried soils of Units 3 and 4 (Fig. 10) contain considerable silt that is likely loess. Pederson et al. (2000) proposed a threepart conceptual model of hillslope responses to climate change, including the suggestion that eolian processes are more active during dryer climate. The buried loess was optically stimulated luminescence dated to estimate the minimum age of the debris flow deposits. Two samples from the soil horizons were collected in May 2005 from the trench exposures to constrain the timing of debris flow deposition and its relation to pluvial climate. Results will be presented during the October 2005 field trip. The early Holocene (Qy1) to recent Holocene alluvial fans are formed by deposition of debris-flow and flood-transported materials from ephemeral streams that drain the escarpment and the western edge of the Uinkaret Plateau. These fans are mostly unconsolidated silts, sands, and gravels, and are quite coarse near the base of the cliffs. The more extensive fans merge with the fine-grained valley fill alluvium seen in Upper and Lower Hurricane Valley. Drainages of the Lower Hurricane Valley flow axially north toward tributaries of the Virgin River. There is considerable incision into Holocene valley fill deposits north of the Grandstand area—some arroyos are more than 10 m deep. This incision is likely related to the post-1860s downcutting (Hereford, 2002). A drainage divide, created by the higher terrain near the Twin Butte salient, isolates the Lower Hurricane Valley from the Upper Hurricane Valley. The Upper Hurricane Valley drainages empty into the Colorado River at the Grand Canyon.

Relict colluvial deposits Faulted deposit

Road Log Return to the main track and turn right (south). Cumulative mi (km)

Figure 12. The Hurricane Escarpment just south of Cottonwood Canyon. Relict colluvial deposits typically found on the middle to upper parts of the escarpment and are a common feature of the steep hillslopes seen along the field trip route. The relict colluvial boulder deposits shown here grade downslope into faulted alluvial fan deposits.

54.3

(87.4)

54.9

(88.4)

Description At 10 o’clock, note the lighter Holocene debris flows overlying faulted Pleistocene fans. Stop 4—Discussion of evidence for Holocene faulting and the salient/section boundary; 30 min.

Stop 4—Evidence of Holocene Faulting? At this stop, the fault scarps observed along the Shivwits section are observed only on late Pleistocene alluvial fans and colluvium. Evidence of fault rupture was sought in the Holocene

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment fan deposits but no definitive scarps or fault-related lineaments were identified. Because much of the Holocene alluvium along the base of the escarpment is dominated by cobble- to-boulder debris-flow deposits with little fine material, small (<0.5 m) surface ruptures might be overlooked. We considered that the modern drainages might preserve evidence of rupturing as knickpoints in the stream-longitudinal profile (Burbank and Anderson, 2001). We measured topographic longitudinal profiles on two modern, sinuous drainages located in T37N, R9W (just east of this location), sections 16 and 21, using an electronic distance measuring total station. One stream has a drainage basin confined to the escarpment, the other stream drains a large re-entrant and has some flow contribution from the Uinkaret Plateau (footwall of the Hurricane Fault). Both stream channels had cobble- to boulder-size clasts with some bedrock exposed along the stream thalweg. In both cases, the stream profiles were nearly linear with no evidence of exposed bedrock or other change in gradient. We concluded that if there had been Holocene surface-rupturing earthquakes along this portion of the Shivwits section, the amount of surface offset was insufficient to be preserved. Cumulative mi (km) 57

(91.8)

57.9

(93.2)

58.8

(94.7)

60

(96.6)

60.7

(97.7)

61.3

(98.7)

Description Twin Butte at 11 o’clock, Diamond Butte at 2 o’clock; note the lower escarpment height around Hurricane Fault salient (Fig. 8), the white Shnabkaib Member of the Moenkopi Formation in Diamond Butte is prominent. Ascending an older Pleistocene (Pliocene?) fan. At 3 o’clock note the recent incision into Holocene and late Pleistocene alluvium likely related to historic arroyo cutting (Hereford, 2002). The windmill and ranch site are the remains of Childer’s Well. We are driving through incised Holocene and Pleistocene alluvium for the next 0.25 mile. Crossing over the pass between Diamond and Twin Buttes; Virgin Limestone Member of the Moenkopi Formation exposed at 3 o’clock in Diamond Butte. At 9 o’clock, note the Hurricane Escarpment and a lower scarp at its base, this small erosional terrace is cut into the easily eroded Moenkopi rocks. The campsite is on the right near the Piñon and Juniper trees, (UTM 12S 288068 4048119). End of Day 2.

Day 3 This section of the road log describes the Whitmore Canyon section of the Hurricane Fault (Fig. 4C). The route roughly parallels the Hurricane Escarpment down to an overlook of the

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Colorado River. Stop 5, mile 74.6 (120.1 km), at the Mount Trumbull schoolhouse is an excellent location to introduce the geology along the Whitmore Canyon section and discuss the style of faulting. At Stop 6, mile 81.9 (131.9 km), the Uinkaret Volcanic Field and its relation to the Hurricane Fault are highlighted. Surface rupture of basalt flows is illustrated at Stop 7, mile 90.3 (145.4 km), located on the faulted Bar Ten basalt flow. Fault morphology on the flow surface and estimates of slip rate in this area are presented. At Stop 8, mile 74.6 (120.1 km), the complexity of faulting of the Whitmore Cascades and slip-rate estimates for the Whitmore Canyon section of the Hurricane Fault are discussed. Stop 9, mile 96.8 (155.9 km), is the southern termination of the field trip. The fault kinematics and geomorphology of lower Whitmore Canyon are discussed; short hikes to examine the evidence of the development of Whitmore Wash are possible from this stop. Road Log Cumulative mi (km) 62.7

(101)

63.1 (101.6)

68.5 (110.3) 69.8 (112.4) 74.6 (120.1)

Description Leave campsite at 8:30 a.m.; turn left onto main road. Turn left at Y intersection toward Mount Trumbull in the distance. There are low hills of Moenkopi Formation for the next several miles. Sharp turns in the road; Hurricane Cliffs at 12 o’clock are framed by ranch house ruins. Shivwits Plateau on the horizon at 3 o’clock. Stop 5—Mount Trumbull schoolhouse.

Stop 5—The Relationship of the Hurricane Fault to the Uinkaret Volcanic Field In the distance to the east, the Hurricane Escarpment is covered by the Pliocene to Holocene basalts of the Uinkaret Volcanic Field (Billingsley et al., 2001). We will discuss the volcanic field in more detail at the next stop. The informally named Basalt of Bundyville (Pliocene; Billingsley et al., 2001) appears to be offset an amount equal to the underlying Moenkopi Formation. This suggests that movement on the Hurricane Fault began here after the emplacement of the basalt at 3.4 Ma (Wenrich et al., 1995). This location is interpreted to be part of the transition from the Shivwits section to the Whitmore section, based on complex discontinuities in the fault trace between here and Whitmore Canyon to the south and a decrease in total fault offset (Stenner and Pearthree, 1999; Billingsley et al., 2001). The schoolhouse is a replica of the historic one-room schoolhouse and town hall that was constructed in 1922, abandoned in the late 1960s and destroyed by vandals in 2000. This area was settled by several Mormon families, the most prominent of which were the Bundys, hence the town’s other name: Bundyville. The

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interior of the schoolhouse contains a variety of historical information and photographs about the settlement and abandonment of this remote area, and is well worth a few minutes exploration. Evidence of Fault Segmentation along the Whitmore Section Like the Shivwits section to the north, the Whitmore section displays relic complexities that suggest a history of fault growth by the linkage of overlapping strands. One such structure is well exposed in lower Whitmore Canyon near the Colorado River. At that location, the main strand of the fault dies out rapidly to the south, whereas another strand to the east rapidly accumulates offset, becoming the principle strand. In the zone between these strands, rocks of the lower Paleozoic section are bent southward, forming a ramp that accommodates the transfer of offset between these strands. Both strands have been active in the late Quaternary based on Raucci (2004), although it is not clear if they are contemporaneously active or if the eastern strand has become the focus of displacement. The structure, displacement history and kinematics of the Hurricane Fault zone in the area will be discussed in detail during Stops 8 and 9. The Hurricane Fault splays significantly near where it crosses the Colorado River, at river mile 191 of the Grand Canyon (the mouth of Whitmore Wash, near field trip Stop 8, is at river mile 188; Fig. 4A). Displacement across the main fault zone approaches zero, while total displacement, ~250 m, is spread across a wide (2 km) zone of discontinuous faults. Because of the complex discontinuity in the trace of the fault and the overall offset minimum, this location is tentatively interpreted as the end of the Whitmore section and the beginning of the newly named Three Springs section (Raucci, 2004). Cumulative mi (km)

Description

77.5 (124.8) Entering the Grand Canyon—Parashant Canyon National Monument. 79.1 (127.4) Mount Emma is seen at 11 o’clock. We are driving over alluvium-covered Paleozoic rocks for the rest of the field trip as we descend toward Whitmore Canyon, we are now on Kaibab Formation carbonates. 79 (127.2) Mount Logan coming into view at 10 o’clock. 81.9 (131.9) Stop 6—Discuss the Uinkaret volcanic field—its age, distribution, and genesis.

basalt flows are more widely distributed than the Pliocene flows and are related to a late eruptive stage that produced many pyroclastic cinder cones (Billingsley et al., 2001) and has continued into the Holocene (Fenton et al., 2001). The rocks of the volcanic field are composed primarily of basalt flows and associated cinder cones, with minor amounts of andesite. Vents and fissures are aligned north-south, such as Mount Emma, Petty Knoll, and Slide Mountain (Billingsley et al., 2001). The eruptive centers are roughly parallel to the Hurricane and Toroweap faults, but are rarely lie along or directly on the faults. The relationship between tectonism and volcanism is not well understood in this area, but they appear to be coeval. The Whitmore section of the Hurricane Fault also shows evidence of Quaternary surface-rupturing events (Fenton et al., 2001; Raucci, 2004; Stenner and Pearthree, 1999), although the inferred displacement rates are lower than on the Shivwits section to the north. Results of morphologic modeling of fault scarps, surface dating and offsets in Pleistocene basalt flows all suggest vertical displacement rates of 0.05–0.11 mm/yr during the late Quaternary. However, these preliminary studies do not tightly constrain the timing and magnitude of the MRE or recurrence intervals. Cumulative mi (km) 81 83.6 86.2

87.1 88.8

90.3

Description

(130.4) Begin the steep decent into upper Whitmore Canyon. (134.6) Now entering Whitmore Canyon. (138.8) Hells Hollow is seen to left, relic colluvial deposits on the steep ridge at 2 o’clock. We are driving on Pleistocene valley fill. Note the large basalt boulders on the alluvium. (140.2) Bar Ten Ranch: a dude ranch and stopover for tourists flying into and out of the Grand Canyon. (143.0) Small airstrip on the right, the Bar Ten basalt flow at 12 o’clock. There are 2–8 m high fault scarps in Pleistocene alluvial fans at 10 o’clock. The Upper Paleozoic section is exposed down to the Esplanade Sandstone in the upper Supai Group. Note the rollover of bedding in the Kaibab Formation on the west wall of the canyon. (145.4) Stop 7 at the Bar Ten basalt flow.

Stop 7—Bar Ten Basalt Flow Stop 6—Uinkaret Volcanic Field The bulk of the Uinkaret Volcanic Field, including its second and third highest peaks, Mount Logan and Mount Emma is visible. Activity began with the eruption of the basalt of Bundyville at 3.6 ± 0.018 Ma (Billingsley et al., 2001). This Pliocene basalt flow caps the soft rocks of the Chinle and Moenkopi Formations at Mount Trumbull and Mount Logan. Pliocene basalt was also deposited along the hanging wall of the Hurricane Fault and to the north of Mount Emma (Billingsley et al., 2001). Quaternary

This stop is atop a significant basalt cascade that issued from vents on the Uinkaret Plateau above and spilled into Whitmore Canyon (Figs. 13 and 14). The Bar Ten flow was emplaced at 88 ± 6 ka based on recent 3He dating (Fenton et al., 2001). The cosmogenic isotope 3He to was used to check previous K-Ar and ArAr dating, which has been problematic in these flows (Dalrymple and Hamblin, 1998; Laughlin et al., 1994). The fault trace is obvious on aerial photos of the area, but a brief look around will demonstrate the difficulty of finding unequivocal fault offsets

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STOP 8

Figure 13. Oblique aerial photo of the Whitmore Canyon area and the Bar Ten and Whitmore Cascade flows. Faults are approximately located. 5

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amidst the complex topography of the basalt flow. Nonetheless, Fenton et al. (2001) suggest an average offset in the flow surface of 10 ± 3 m based on interpretation of several long topographic profiles (Stenner and Pearthree, 1999), for a time-averaged slip rate of ~110 m/m.y. (0.11 mm/yr). This is less than estimates for the Shivwits section (Amoroso et al., 2004) but at the high end of estimates from the Whitmore section (Fenton et al., 2001; Raucci, 2004), suggesting a southerly attenuation of slip along the Hurricane Fault. Some of the regional deformation budget may be accommodated at this latitude by the Toroweap Fault to the east, which becomes more active to the south (Jackson, 1990).

36 10' N

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ra Colo 191

9 Gray Ledge

Description

90.9 (146.4) A thin wedge of Coconino Sandstone outcrops on the Hermit Formation. The Coconino Sandstone thickens to over 200 m in the eastern Grand Canyon. Note the elongated deposits of relict boulder colluvium on the west wall of canyon. Whitmore Cone (a large dark cinder cone) comes into view at 10 o’clock. 92.3 (148.6) Thick Pleistocene debris flows at 2 to 4 o’clock. 92.8 (149.1) A significant fault scarp in alluvium to the left of the road is visible for next several hundred meters. Topographic profiles (Stenner and Pearthree, 1999) showed a displacement range of 4.1–12 m. Ages of two displaced alluvial fans were estimated based on 3He cosmogenic age estimates (Fenton et al.,

192

192 Mile Canyon

188 River miles Fault, dashed where approximately located Quaternary basalt flows Quaternary alluvial deposits Cinder cone

Figure 14. Simplified geologic map of the Hurricane Fault in lower Whitmore Canyon, and along the Colorado River from mile 188 to mile 192 (river miles downstream of Lee’s Ferry). Field trip stops are shown. Numbers refer to fault scarps and sampling locations (Table 1).

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2001). Using the mean alluvial fan age and mean displacements resulted in slip rate estimates of 0.11–0.28 mm/yr. This area was also used for the carbonate rind thickness-age calibration (Amoroso et al., 2004). The Whitmore Cascade flow is ahead. 93.3 (150.2) An abrupt step in Whitmore Cascade basalt probably correlative with linear scarp noted above. Displacement of the basalt was estimated to be 11.4–17.6 m and the surface age of Whitmore Cascade was 140–180 ka (mean age, 153 ka; Fenton et al., 2001) so the slip rate is 0.05–0.13 mm/yr. The vegetation has changed considerably from the sagebrush and yucca of the Shivwits section because of the lower elevation. We are now in the lower Sonoran life zone. 95.6 (153.9) Stop 8—Displaced basalts of lower Whitmore Canyon (UTM 12S 301426 4005310). Stop 8—Faulting of the Whitmore Cascade Flow This is the Whitmore Cascade flow, which also originated atop the Uinkaret Plateau (Fig. 13), around 180 ka (Fenton et al., 2001; Raucci, 2004). At this stop evidence of mid-late Pleistocene activity on the Hurricane Fault can be seen. The Hurricane Fault probably consists of two strands here, one of which lies under the road with the other lying several hundred meters to the west (Fig. 14). The geometry of the fault zone is unfortunately hidden by the basalt flow. Stenner and Pearthree (1999), Fenton et al. (2001), and Raucci (2004) all reported offsets of 12–15 m in this flow and offset rates of 0.07–0.08 mm/yr across the strand that underlies the

road (Table 1; Fig. 14). A short hike will bring us to the western strand, where Raucci (2004) reported a 13.8 m offset in an older basalt flow (540 ± 30 ka) that is partially overlain by the Whitmore Cascade flow (Fig. 14). This fault strand probably accommodated the bulk of Paleozoic bedrock offset, but the observed displacement in the basalt yields an average rate of 0.025 mm/yr. Fault slip may have transferred to a relatively new, eastern strand during the time between emplacement of the flows, or the strands may be active contemporaneously, yielding a combined rate of ~0.09–0.1 mm/yr across the entire zone. The fault zone itself is well exposed here, and the downcutting by Whitmore Wash allows an opportunity to discuss the interaction of repeated basalt flows, stream erosion, and active faulting at this location (see Fig. 15). Road Log Cumulative mi (km)

Description

95.6 (153.9) Begin steep descent of the Whitmore Cascade lava flow. 96.7 (155.7) Note the basalt remnant high atop a hill of Hermit Formation across the Whitmore Canyon. The 3.4 Ma basalt protects a wedge of Coconino and Kaibab Formation derived colluvium, suggesting that the broad platform atop the Esplanade Sandstone had not formed at this time and the canyon rim was much closer to the river. 96.8 (155.9) Stop 9—Fault kinematics and geomorphology of lower Whitmore Canyon. This area is our final stop.

TABLE 1. FLOW AGE AND DATING METHOD, AMOUNT OF VERTICAL OFFSET, AND SLIP RATE FOR BASALT FLOWS AND ALLUVIAL FANS ALONG THE WHITMORE CANYON SECTION OF THE HURRICANE FAULT Location* 1 2 3 4 5 6 7 8 9

Surface Bar Ten flow† alluvial fan† alluvial fan† Whitmore Cascade flow† older Whitmore flow§ Whitmore Cascade flow† Whitmore Cascade flow§ alluvial fan† Gray Ledge flow

Age (ka) 88 ± 6 29 ± 9 74 ± 16 177 ± 9 540 ± 30 177 ± 9 185 ± 26 51 ± 9 97 ± 32**

Method 3

He He 3 He 3 He 39 Ar/40Ar 3 He 39 Ar/40Ar 3 He 39 Ar/40Ar 3

Offset (m)

Rate (mm/yr)

10 ± 3 3 7 15 ± 3 13.8# 12 13.1 4†† ~5#

0.11 0.10 0.09 0.08 0.024–0.027 0.07 0.062–0.082 0.08 ~0.05

*Refer to Figure 14 for locations. † Data from Fenton et al. (2001); 3He ages are based on accumulated cosmogenic. 3He due to exposure and should be considered minimum ages. Errors are ±1σ. § Raucci (2004). # Due to multiple fault strands, these scarps may not capture all of the motion since the emplacement of the basalt flows. **(Pederson et al., 2002). †† Offset estimated using a projection of the far-field slopes (Fenton et al., 2001).

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Stop 9—Lower Whitmore Canyon We have descended the Whitmore Cascade flow and stopped here on the brink of the cliffs above the Colorado River. Under the road, the eastern fault strand has grown to become the primary fault strand, whereas the western strand, which we visited at the last stop, has diminished. The eastern strand under our feet accommodates the bulk of the bedrock offset and cuts the Whitmore Cascade, creating an offset of 13.1 m (Table 1, Fig. 16). The transfer of offset between these fault strands is accomplished by the formation of a relay ramp between the two faults (Figs. 17A and 17B), which is well exposed to the north of our stop. Looking into the gorge of lower Whitmore Canyon, note the thickness of the Whitmore Cascade flow—it is likely that the old course of Whitmore Wash was under the Whitmore Cascade flow, and it has since re-incised a path along the fault to the west. One could make several hikes from this location to examine aspects of the Whitmore Cascade flow. A 2 mi (3.2 km, roundtrip) hike down the Whitmore Trail to the Colorado River is a good way to examine the Whitmore Cascades flow and the lower Paleozoic stratigraphy. This historic trail was once used by the Bar Ten ranch as a re-supply route for river runners, and offers spectacular views of the river corridor and columnar jointing in intra-canyon basalt flows. Another possibility is to hike up Whitmore Canyon to observe the damage zone of the Hurricane Fault and look at kinematic indicators. Raucci (2004) studies along this section of the fault have tentatively suggested SW-oriented extension across this section of the fault. This is in contrast to the EW-oriented extension noted on other sections of the Hurricane Fault, and implies some component of left-lateral oblique slip.

D

t

Older Whitmore flow

Mississippian Redwall Limestone.

Figure 15. Looking south at displacement of the Whitmore Cascade flow by a splay of the Hurricane Fault. The total vertical displacement (black arrow) was measured by projecting the flow surfaces into the fault.

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Stop #8

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This field trip is a brief introduction to the geomorphology and neotectonism of the Shivwits and Whitmore Canyon sections of the Hurricane Fault. While some understanding of the seismic hazard posed by the Hurricane Fault has been gained, more work needs to be done to understand the rupture history of the various sections, identify seismogenic rupture barriers, and the relation of the Hurricane Fault to nearby faults. Much of the Shivwits section is dominated by the Hurricane Escarpment, created as the result of vertical displacement along the Hurricane Fault. The style of alluvial deposition along the fault as the result of erosion or mass movement of hillslope materials has been affected by changes in Pleistocene climate. The present day hillslopes have only a thin mantle of transportable materials, much of the colluvial mantle may have been removed during the late Pleistocene. Neotectonism along the Hurricane Fault has implications for both the seismic hazard as well as landscape development. Paleoseismic investigations on the Shivwits section have estimated that the slip rate since the late Quaternary has been ~0.1–0.3 mm/yr. Trench studies have revealed that this portion of the fault has the potential for M7 earthquakes that would affect the populated

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13.8 m vertical displacement

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Figure 16. View northeast across lower Whitmore Canyon, toward Stop 9.

southwestern Utah area. Evidence was presented that suggests these sections of the Hurricane Fault are the result of the linkage of smaller fault strands. The re-entrant area of the Shivwits section is a good example of the linkage of fault strands and how strain is accommodated along the fault. Fault linkages have a strong control on the drainage systems that cross the escarpment.

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Figure 17. (A) Simplified tectonic map of the Hurricane Fault zone in lower Whitmore Wash, showing major faults and location of fault relay. All faults shown are down to the west normal faults. The pointing hand shows the viewing direction of the photograph in Figure 15. (B) Cross section of the Hurricane Fault zone ~300 m north of Stop 9. (C) Cross section of Hurricane Fault zone, ~200 m south of Stop 9.

The Whitmore Canyon section physiography is considerably different from the Shivwits section because of the Uinkaret Volcanic Field is astride the Hurricane Fault. The Whitmore Canyon section has been shown to be less active than the Shivwits section. Studies of faulted basalt flows and alluvial fans show that the 0.05–0.10 mm/yr slip rate of this section is lower than the Shivwits section. The change of course of the Whitmore Wash in the late Pleistocene is more evidence of the strong influence of tectonism on the geomorphology along the Hurricane Fault. Neotectonics of Subsidiary Faults Related to the Hurricane Fault The Washington, Mokaac, Main Street, and Sunshine faults (Billingsley and Workman, 2000) are small, active

normal faults, west of the Shivwits section re-entrant, seen on the field trip. This pattern of faulting is probably the result of compressional Laramide tectonics later reactivated by Late Cenozoic extension as suggested for the Grand Wash, Hurricane, Toroweap-Sevier, and Kaibab faults. What is the tectonic relationship between the Hurricane Fault and these faults? There is considerable evidence that these faults show late Quaternary activity but no paleoseismic investigation to estimate slip rate or age of last rupture has been done. Can rupture of the Hurricane Fault be transferred to these smaller faults? Another though shallower re-entrant is seen in the Whitmore section of the Hurricane Fault. There is a similar pattern of short, normal faults west of the Hurricane Fault, along the Whitmore section, that include the Dellenbaugh, Frog, Grassy, and Andrus faults (Billingsley and Wellmeyer,

Paleoseismology and geomorphology of the Hurricane Fault and Escarpment 2003). These faults may have an analogous relationship as those along the Shivwits section. Fault Pattern Complexity of the Hurricane Fault Zone The fault pattern complexity discussed above is also seen at re-entrants along the Lemhi and Lost River normal faults in eastcentral Idaho (Janecke, 1993). Is the complexity related to the geometry of segmented normal faults (Peacock and Sanderson, 1991; Willemse, 1994)? If the Hurricane Fault is listric at depth (Hamblin, 1965), and the fault strike changes from N15W near the Anderson Junction–Shivwits boundary through N25W at the Grandstand to N55E north of Twin Buttes, does this form a nonconservative geometric barrier (King, 1986; King and Yielding, 1984) to rupture? Put another way, rupture propagating into curving fault may result in accommodation space problems in a zone of convergent slip; the resulting fragmentation barrier may stop rupture propagation. Do the Washington, Mokaac, Main Street, and Sunshine faults extend to the seismogenic zone and meet a possible lowangle portion of the Hurricane Fault at depth (Erskine, 2001)? If so, ruptures that initiate at the base of the seismogenic zone may propagate upward as well as along the low-angle portion of the Hurricane Fault (Ofoegbu and Ferrill, 1998). See Abers (1991), Jackson and White (1989), and Wernicke (1995) for discussions of low-angle faults and seismicity. The seismic hazard potential presented by the Hurricane Fault should include assessments of these smaller faults. ACKNOWLEDGMENTS We thank Phil Pearthree (Arizona Geological Survey) and Sarah Robinson (U.S. Geological Survey) for their thoughtful reviews and Tracey Felger (U.S. Geological Survey) for help with the road log maps. REFERENCES CITED Abers, G.A., 1991, Possible seismogenic shallow-dipping normal faults in the Woodlark-D’Entrecasteaux extensional province, Papua New Guinea: Geology, v. 19, p. 1205–1208, doi: 10.1130/0091-7613(1991)019<1205: PSSDNF>2.3.CO;2. Amoroso, L., 2001, Studies in Quaternary Geology of Arizona: Active Tectonics—Relationship of Soils to Surficial Geology [Ph.D. dissertation]: Tempe, Arizona State University, 260 p. Amoroso, L., 2005, Age calibration of carbonate rind thickness in late Pleistocene soils for surficial age estimation, southwest USA: Quaternary Research (in press), doi: 10.1016/j.yqres.2005.06.003. Amoroso, L., Pearthree, P.A., and Arrowsmith, J.R., 2004, Paleoseismology and Neotectonics of the Shivwits Section of the Hurricane Fault, Northwestern Arizona: Bulletin of the Seismological Society of America, v. 94, p. 1919–1942, doi: 10.1785/012003241. Anderson, R.E., and Mehnert, H., H., 1979, Reinterpretation of the history of the Hurricane Fault in Utah, in Newman, G.W., and Goode, H.D., eds., 1979 Basin and Range Symposium: Denver, Colorado, Rocky Mountain Association of Geologists, p. 145–165. Arabasz, W.J., and Julander, D.R., 1986, Geometry of seismically active faults and crustal deformation within the Basin and Range–Colorado Plateau transition in Utah, in Mayer, L., ed., Extensional tectonics of the south-

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western United States: A perspective on processes and kinematics: Geological Society of America Special Paper 208, p. 43–74. Arabasz, W.J., Pechmann, J.C., and Brown, E.D., 1992, Observational seismology and the evaluation of earthquake hazards and risk in the Wasatch Front area, Utah, in Gori, P., and Hays, W.W., eds., Assessment of the Regional Earthquake Hazards and Risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1500-A-J, p. D1–D36. Billingsley, G.H., 1991, Geologic map of the Sullivan Draw North quadrangle, northern Mohave County, Arizona: U.S. Geological Survey Open-File Report 91-558. Billingsley, G.H., 1993a, Geologic map of the Dutchman Draw quadrangle, northern Mohave County, Arizona: U.S. Geological Survey Open-File Report 93-587. Billingsley, G.H., 1993b, Geologic map of the Grandstand quadrangle, northern Mohave County, Arizona: U.S. Geological Survey Open-File Report 93-588. Billingsley, G.H., 1993c, Geologic map of the Russell Spring quadrangle, northern Mohave County, Arizona: U.S. Geological Survey Open-File Report 93-717. Billingsley, G.H., 1993d, Geologic map of the Wolf Hole Mountain and vicinity, Mohave County, northwestern Arizona: U.S. Geological Survey Map I-2296, scale1:31,680. Billingsley, G.H., 1994a, Geologic map of the Antelope Knoll quadrangle, northern Mohave County, Arizona: U.S. Geological Survey Open-File Report 94-449. Billingsley, G.H., 1994b, Geologic map of the Moriah Knoll quadrangle, northern Mohave County, Arizona: U.S. Geological Survey Open-File Report 94-634. Billingsley, G.H., and Wellmeyer, J.L., 2003, Geologic map of the Mount Trumbull 30′ × 60′ quadrangle, Mohave and Coconino Counties, northwestern Arizona: U.S. Geological Survey Geologic Investigations Series I-2766, 36 p. Billingsley, G.H., and Workman, J.B., 2000, Geologic map of the Littlefield 30′ × 60′ quadrangle, Mohave County, northwestern Arizona: U.S. Geological Survey Geologic Investigations Series I-2628, 25 p. Billingsley, G.H., Hamblin, W.K., Wellmeyer, J.L., and Dudash, S.L., 2001, Geologic map of part of the Uinkaret Volcanic Field, Mohave County, northwestern Arizona, U.S. Geological Survey Map MF-2368, scale 1:31,680. Bull, W.B., 1991, Geomorphic responses to climatic change: New York, Oxford University Press, 326 p. Burbank, D.W., and Anderson, R.S., 2001, Tectonic Geomorphology: Malden, Massachusetts, Blackwell Science, 274 p. Christensen, G.E., 1995, The September 2, 1992, ML 5.8 St. George earthquake: Utah Geological Survey Circular, 41 p. Crone, A.J., and Haller, K.M., 1991, Segmentation and the coseismic behavior of Basin and Range normal faults: examples from east-central Idaho and southwestern Montana, U.S.A.: Journal of Structural Geology, v. 13, p. 151–164, doi: 10.1016/0191-8141(91)90063-O. Dalrymple, G.B., and Hamblin, W.K., 1998, K-Ar of Pleistocene lava dams in the Grand Canyon in Arizona: Proceedings of the National Academies of Science of the United States of America, v. 95, p. 9744–9749, doi: 10.1073/pnas.95.17.9744. dePolo, C.M., Clark, D.G., Slemmons, D.B., and Ramelli, A.R., 1991, Historical surface faulting in the Basin and Range province, western North America: implications for fault segmentation: Journal of Structural Geology, v. 13, p. 123–136, doi: 10.1016/0191-8141(91)90061-M. Doser, D.I., 1985, The 1983 Borah Peak, Idaho and 1959 Hebgen Lake, Montana earthquakes: Models for normal fault earthquakes in the Intermountain Seismic Belt, in Stein, R.S., and Bucknam, R.C., eds., Proceedings Workshop XXVIII, on the Borah Peak, Idaho Earthquake: U.S. Geological Survey Open-File Report 85-290, v. A, p. 368–384. Downing, R.F., Smith, E.I., Orndorff, R.I., Spell, T.L., and Zanetti, K.A., 2001, Imaging the Colorado Plateau-Basin and Range transition zone using basalt geochemistry, geochronology, and geographic information systems, in Erskine, M.C., Faulds, J.E., Bartley, J.M., and Rowley, P.D., eds., The Geologic Transition, High Plateaus to Great Basin—A Symposium and Field Guide: Utah Geological Association Publication 30, p. 127–154. Erskine, M.C., 2001, Colorado Plateau tectonostratigraphic unit, in Erskine, M.C., Faulds, J.E., Bartley, J.M., and Rowley, P.D., eds., The Geologic Transition, High Plateaus to Great Basin—A Symposium and Field Guide, Volume 30: Salt Lake City, Utah Geological Association, p. 39–56. Fenton, C.R., Webb, R.H., Pearthree, P.A., Cerling, T.E., and Poreda, R.J., 2001, Displacement rates on the Toroweap and Hurricane Faults: Implications for Quaternary downcutting in the Grand Canyon, Arizona:

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McCalpin, J.P., and Nelson, A.R., 1996, Introduction to paleoseismology, in McCalpin, J.P., ed., Paleoseismology: San Diego, California, Academic Press, p. 1–32. Menges, C.M., and Pearthree, P.A., 1983, Map of neotectonic (Latest PlioceneQuaternary) deformation in Arizona, Arizona Department of Geology and Mineral Technology, p. 48, 4 sheets, scale 1:500,000, 1:133,830, and 1:121,000. Ofoegbu, G.I., and Ferrill, D.A., 1998, Mechanical analysis of listric faulting with emphasis on seismicity assessment: Tectonophysics, v. 284, p. 65–77, doi: 10.1016/S0040-1951(97)00168-6. Ostenaa, D., 1984, Relationships affecting estimates of surface fault displacements based on scarp-derived colluvium deposits: Geological Society of America Abstracts with Programs, v. 16, p. 327. Peacock, D.C.P., and Sanderson, D.J., 1991, Displacements, segment linkage and relay ramps in normal fault zones: Journal of Structural Geology, v. 13, p. 721–733, doi: 10.1016/0191-8141(91)90033-F. Peacock, D.C.P., and Sanderson, D.J., 1994, Geometry and development of relay ramps in normal fault systems: American Association of Petroleum Geologists Bulletin, v. 78, p. 147–165. Pearthree, P.A., 1998, Quaternary fault data and map of Arizona: Arizona Geological Survey, p. 122, 1 sheet, scale 1:750,000. Pearthree, P.A., and Bausch, D.B., 1999, Earthquake hazards in Arizona, Arizona Geological Survey Map 34, scale 1:1,000,000. Pederson, J., Pazzaglia, F., and Smith, G., 2000, Ancient hillslope deposits: Missing links in the study of climate control on sedimentation: Geology, v. 28, p. 27–30, doi: 10.1130/0091-7613(2000)028<0027:AHDMLI>2.3.CO;2. Pederson, J., Karlstrom, K., Sharp, W., and McIntosh, W., 2002, Differential incision of the Grand Canyon related to Quaternary faulting—Constraints from U-Series and Ar/Ar dating: Geology, v. 30, p. 739–742, doi: 10.1130/0091-7613(2002)030<0739:DIOTGC>2.0.CO;2. Powell, J.W., 1875, Exploration of the Colorado River of the West and its Tributaries, 1869–1872: Washington, D.C., Government Printing Office, 213 p. Raucci, J.J., 2004, Structure and neotectonics of the Hurricane Fault, western Grand Canyon, Mojave County, Arizona [M.S. thesis]: Flagstaff, Arizona, Northern Arizona University, 188 p. Reynolds, S.J., Florence, F.P., Welty, J.W., Roddy, M.S., Currier, D.A., Anderson, A.V., and Keith, S.B., 1986, Compilation of radiometric age determinations in Arizona: Tucson, Arizona Bureau of Geology and Mineral Technology Bulletin 197, 2 sheets, scale 1:1,000,000, p. 258. Scarborough, R.B., Menges, C.M., and Pearthree, P.A., 1986, Late PlioceneQuaternary (post 4 m.y.) faults, folds, and volcanic rocks in Arizona: Arizona Bureau of Geology and Mineral Technology, Map 22, scale 1: 1,000,000. Schlische, R.W., 1995, Geometry and origin of fault-related folds in extensional settings: American Association of Petroleum Geologists Bulletin, v. 79, p. 1661–1678. Schwartz, D.P., 1988, Geologic characterization of seismic sources: Moving into the 1990’s, in Thun, J.L.V., ed., Earthquake Engineering and Soil Dynamics II—Recent Advances in Ground-motion Evaluation: New York, American Society of Civil Engineers Geotechnical Special Publication 20, p. 1–42. Schwartz, D.P., and Coppersmith, K.J., 1984, Fault behavior and characteristic earthquakes: Examples from the Wasatch and San Andreas fault zones: Journal of Geophysical Research, v. 89, p. 5681–5698. Smith, R.B., and Arabasz, W.J., 1991, Seismicity of the Intermountain seismic belt, in Slemmons, D.B., Engdahl, E.R., Zoback, M.D., and Blackwell, D.D., eds., Neotectonics of North America: Geological Society of America, Geology of North America, p. 185–228. Sorauf, J.E., and Billingsley, G.H., 1991, Members of the Toroweap and Kaibab Formations, lower Permian, northern Arizona and southwestern Utah: The Mountain Geologist, v. 28, p. 9–24. Stenner, H.D., and Pearthree, P.A., 1999, Paleoseismology of the southern Anderson Junction section of the Hurricane Fault, northwestern Arizona and southwestern Utah, in Stenner, H.D., Lund, W.R., Pearthree, P.A., and Everitt, B.L., eds., Paleoseismic Investigations of the Hurricane Fault in Northwestern Arizona and Southwestern Utah: Tucson, Arizona Geological Survey Open-File Report 99-8, p. 45–81. Spencer, J.E., and Reynolds, S.J., 1989, Middle Tertiary tectonics of Arizona and adjacent areas, in Jenney, J.P., and Reynolds, S.J., eds., Geologic Evolution of Arizona: Tucson, Arizona Geological Society Digest 17, p. 539–574. Stenner, H.D., Lund, W.R., Pearthree, P.A., and Everett, B.L., 1998, Quaternary history and rupture characteristics of the Hurricane Fault, southwestern

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Printed in the USA

Geological Society of America Field Guide 6 2005

Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior, Book Cliffs, Utah Simon A.J. Pattison Department of Geology, Brandon University, Brandon, Manitoba R7A 6A9, Canada

ABSTRACT Marine mudstone-encased, inner shelf sandstone bodies are concentrated in a 70–90-m-thick interval that spans the upper Aberdeen and lower Kenilworth members (Blackhawk Formation, Campanian), Book Cliffs, eastern Utah. These sandstone bodies contain a complex mixture of event beds including wave or storm-modified turbidites, hummocky cross stratified sandstones, hyperpycnites, and/or classical turbidites. The turbiditic channel-fills and lobes were deposited below fair weather wave base and are detached from their time equivalent shoreface deposits. Shallow marine facies models should be revised to include turbiditic-rich channels and lobes in some inner shelf settings. A three-component shoreface-to-shelf model, consisting of delta front deposits, subaqueous channels, and prodelta turbidites, is proposed to explain the depositional setting and environment of the Mancos Shale–encased sandstone bodies. Oceanic- or river-flood induced hyperpycnal flows were responsible for cutting a network of subaqueous channels on the inner shelf and for transporting fine-grained sediments from the shoreface to the inner shelf. Other Mancos Shale– encased isolated sandstone bodies in eastern Utah and western Colorado should be reexamined in the light of the new data and models presented herein. Keywords: inner shelf, turbidites, storms, Book Cliffs, Western Interior Seaway, sedimentology, sequence stratigraphy. INTRODUCTION Isolated sandstone bodies of enigmatic origin are observed throughout the Mancos Shale of eastern Utah and western Colorado. Many of these bodies have a channel-form or lobate geometry and are dominated by turbidite-like event beds. A variety of interpretations have been proposed over the past 25 years, including lowstand shoreface and valley-fill systems, distributary mouth bars and distributary channels, and prodelta plumes and shelf deposits (Balsley, 1980; Swift et al., 1987; Chan et al., 1991; Cole and Young, 1991; Chan, 1992; O’Byrne and Flint, 1995; Cole et al., 1997; Hampson et al., 1999). The main purpose of this field trip is to scrutinize the sedimentology, sedimentary

architecture, depositional models, up-dip correlation, sequence stratigraphy, and paleogeography of the Mancos Shale–encased, isolated sandstone bodies in the Green River to Thompson area, Book Cliffs, eastern Utah. These bodies are concentrated in a relatively narrow stratigraphic interval (70–90 m thick) and are dominated by turbiditic-rich channel fills and lobes. Recent research has revealed a complex mixture of wave or storm-modified turbidites, hummocky cross stratified sandstones, hyperpycnites, or classical turbidites in these bodies. Hyperpycnal flows are gaining increased acceptance as an important transporting agent of silt and fine sand from the shallow marine to inner shelf environments. Both storm-induced underflows (oceanic floods) and river flood underflows have

Pattison, S.A.J., 2005, Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior, Book Cliffs, Utah, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 479–504, doi: 10.1130/2005.fld006(21). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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Figure 1. Book Cliffs of eastern Utah showing the location of the field trip stops, major highways, and rivers. Gray shaded rectangles highlight the detailed field maps (Figs. 3, 5, 7, and 9). Inset maps show the position of the study area. PRC—Price River Canyon; ME—Mount Elliot; DS—Desert Siding; BB—Battleship Butte; MM—Middle Mountain; GB—Gunnison Butte; TC—Tusher Canyon; HM—Hatch Mesa; FW— Floy Wash; BW—Bootlegger Wash; SW—Sagers Wash; SLC—Salt Lake City.

been well documented on modern shelves (Wright et al., 1988; Wheatcroft, 2000); however, the resulting deposits are rarely described from the rock record. Their paucity in the rock record can be explained in one of two ways: (1) underflows were not an important component of shoreface-shelf systems in the past, or (2) underflow deposits have been overlooked or misidentified. The balance of the evidence presented on this field trip supports the latter explanation. Shoreline to shelf facies models and sequence stratigraphic models are currently undergoing revision to show the depositional setting and stratigraphic position of inner shelf turbidites bodies (Pattison, 2005a, 2005b). Mancos Shale–encased isolated sandstone bodies in eastern Utah and western Colorado should be reexamined in the light of the new models presented herein. GEOLOGICAL SETTING The Books Cliffs arguably represent the best exposed and studied deltaic rocks in the world. These famous rocks have been

used to develop, test, and refine sedimentological and stratigraphic ideas and models over the years, including the principles and concepts of sequence stratigraphy (e.g., Van Wagoner et al., 1990). The Book Cliffs extend for ~300 km from Helper, Utah, into western Colorado. The southern sector of these cliffs, located in eastern Utah, is the focus for this field trip (Fig. 1). The Book Cliffs consist of Campanian (Upper Cretaceous) siliciclastic rocks that were sourced from the Sevier highlands to the west and were deposited as a series of north-south–trending depositional belts to the east. Each rock package consists of fluvial and coastal plain deposits in the west and time equivalent shoreface to shelf deposits in the east. During the Campanian, the Cretaceous Western Interior Seaway covered the eastern half of Utah, with paleoshoreline trends oriented approximately north to south (Tables 1 and 2) (McGookey et al., 1972). The Book Cliffs in east-central Utah consist of the Star Point Formation, Blackhawk Formation, and Castlegate Sandstone, all of which are part of the Mesaverde Group (Young, 1955). The Blackhawk Formation is subdivided into the Spring Canyon,

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TABLE 1. SEDIMENT TRANSPORT DIRECTION Stratigraphic unit

Paleocurrent direction (°E)

n

Lower Kenilworth Member (KPS2)

N72 N65 N101 N75 N92 N72 N73 N71

746 5 13 23 34 125 62 484

Top Aberdeen Member (CG Layer)

N145 N141 N134 N159

53 33 6 14

Upper Aberdeen Member (Level 2)

N109 N118 N80 N107 N82

Upper Aberdeen Member (Level 1)

N107 N127 N69 N111 N66 N123 N101

Locations

References

Cumulative/weighted mean Middle Mountain Toad Stool Bootlegger Wash Hatch Mesa sole marks, current ripples Hatch Mesa current ripples Hatch Mesa current ripples, sole marks Lower Sequences and Blue Castle Butte Hatch Mesa current ripples, sole marks

This study This study Hampson et al. (1999) Taylor and Lovell (1991, 1995) Swift et al. (1987) Swift et al. (1987) Newman (1985); Chan et al. (1991)

Cumulative/weighted mean Floy Wash Middle Mountain Toad Stool Floy Wash

This study This study Chan et al. (1991)

164 94 13 42 15

Cumulative/weighted mean Gunnison Butte channel-fill current ripples Gunnison Butte channel axes X-Interval channels X-Interval sheets

This study This study Hampson et al. (1999) Hampson et al. (1999)

179 53 5 44 37 36 4

Cumulative/weighted mean Gunnison Butte channel-fill current ripples Gunnison Butte channel axes Y-Interval channels Y-Interval sheets Tusher Canyon channel-fill current ripples Tusher Canyon channel axes

This study This study Hampson et al. (1999) Hampson et al. (1999) Chan et al. (1991) Chan et al. (1991)

Note: n—number.

Aberdeen, Kenilworth, Sunnyside, Grassy, and Desert members (Fig. 2). These units grade from sandstone-dominated intervals in the west to the mudstone-dominated Mancos Shale in the east (Young, 1955; Balsley, 1980). However, some parts of the Mancos Shale in eastern Utah and western Colorado show an increase in sandstone content (Kellogg, 1977; Cole and Young, 1991). These deposits were originally part of the Mancos B interval and are now formally defined as the Prairie Canyon Member (Fig. 2) (Cole et al., 1997). The lower Kenilworth Member, upper Aberdeen Member, and Prairie Canyon Member are the focus of this field trip. FIELD TRIP ITINERARY This three-day field trip is organized to start in the distal offshore and move progressively landward during the duration of the trip. A wide variety of Mancos Shale–encased sandstone bodies will be examined in the Bootlegger Wash area during Day 1, including thin-bedded turbidite sheets, ironstones, and sandstone- and mudstone-rich turbidite channel-fill packages. There are four stops in the Bootlegger Wash area during the first day (Fig. 1). Day 2 will begin with a 10-mi drive to the Hatch Mesa area to look at the storm-influenced prodelta turbidite complex in the lower Kenilworth Member. Proximal and distal facies will be examined during the four stops at Hatch Mesa (Fig. 1). The final two stops of Day 2 will be north of Green River at the entrance

TABLE 2. PALEOSHORELINE TREND Stratigraphic unit

Paleoshoreline orientation

References

Kenilworth parasequence 2 (KPS2)

N18°W

Taylor and Lovell (1991, 1995); Pattison (1994a)

Kenilworth parasequence 1 (KPS1)

N18°W

Taylor and Lovell (1991, 1995); Pattison (1994a)

Aberdeen parasequence 5 (APS5)

N14°E

Howell and Flint (2003)

Aberdeen parasequence 4 (APS4)

N12°E

Howell and Flint (2003)

Aberdeen parasequence 3 (APS3)

N16°E

Howell and Flint (2003)

to Tusher Canyon to view turbidite-rich channel deposits in the upper Aberdeen Member (Fig. 1). Day 3 will begin with a short drive through the Willow Bend ranch to examine the Mancos Shale–encased channel-fill deposits in the upper Aberdeen Member and a spectacular exposure of lower Kenilworth Member turbidites on the eastern side of Gunnison Butte (Fig. 1). The final three stops of Day 3 will be in Price River Canyon to compare and contrast the lower Kenilworth turbidites with the classic shoreface deposits in the upper Kenilworth Member (Fig. 1).

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482

Figure 2. Scaled-stratigraphic section of the Castlegate Sandstone, Blackhawk Formation, Star Point Formation, and Mancos Shale of east-central Utah (modified from Young, 1955; Balsley, 1980; Cole et al., 1997). Cross section is oriented from west to east along depositional dip and covers the area from Helper, Utah, to the Utah-Colorado border. Distance and thickness data are compiled from measured sections, photo-panoramas, well logs, outcrop correlations, and subsurface correlations (Pattison, 2005a). Note the occurrence of isolated shelf sandstone bodies in the upper Aberdeen Member to lower Kenilworth Member stratigraphic interval. Location of field trip areas are marked by arrows and abbreviations along the top of the stratigraphic section. Abbreviations as per Figure 1. BH CP—Blackhawk Formation coastal plain; C. Ss.—Castlegate Sandstone; Mst—Mudstone.

DAY 1 TRIP LOG: BOOTLEGGER WASH–SAGERS WASH

35.6

(57.3)

Depart the Salt Palace Convention Center in Salt Lake City. Travel south on I-15 for ~50 mi. Take Exit 261 (east) into Spanish Fork and travel east along Hwy 6. Continue east on Hwy 6 via Soldier Summit, Price, Wellington and Woodside until the junction with I-70, which is ~120 mi southeast of Spanish Fork. Turn eastbound onto I-70 (Exit 156) and begin your mileage log.

36.8

(59.2)

38.1

(61.3)

38.8

(62.4)

39.0

(62.8)

Cumulative mi (km) 0.0 2.0 24.0

(0.0) (3.2) (38.6)

29.0

(46.7)

34.0 34.4

(54.7) (55.3)

35.1

(56.5)

Description Exit 156 on I-70. Drive east on I-70. Pass Exit 158 for Green River. Continue east on I-70. Continue east on I-70 past Exit 180 at Crescent Junction (junction with Hwy 191 to Moab). Continue east on I-70 past Exit 185 at Thompson. Take Ranch Exit 190. Turn right off I-70. Turn left and travel north. Bridge over I-70. “Paved” road curves northwest and is deeply rutted. Turn right (E) at the T-junction and drive east along the old highway.

Turn north onto the Sagers Canyon dirt road (note the BLM sign for Sagers Canyon). Enter Sagers Wash and pass underneath a railway bridge. Turn left (NW) onto a jeep track and travel NW-W. Be aware of deep pot holes and road wash-outs. Drive past a vegetated dried-out dugout reservoir on the south side of the road. Park vehicles and disembark. Hike ~0.2 mi across the desert floor toward the southwest. Climb a small hill.

Stop 1.1—Bootlegger Wash–Sagers Wash: Overview Bootlegger Wash and Sagers Wash are located ~45 km east of Green River (Figs. 1 and 3). The desert floor topography in this area is characterized by numerous small hills and ridges, 10– 50 m high, that are capped by interbedded mudstones, siltstones, and very fine-grained sandstones of the Prairie Canyon Member (Cole et al., 1997; Hampson et al., 1999). The Book Cliffs are located ~2 km north and are 150–250 m high in this area (Figs. 1 and 3). The uppermost cliff-forming package is 55–70 m thick and consists of channel and upper shoreface sandstones of the Castlegate Sandstone and Desert Member (Van Wagoner, 1991, 1995; Pattison, 1994b; Nummedal et al., 2001). Marine mud-

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Figure 3. Detailed field map of the Salt Wash, Bootlegger Wash, and Sagers Wash area. Location highlighted in Figure 1. Outcrop of the top Castlegate Sandstone, lower Kenilworth Member distal turbidite lobe (i.e., Bootlegger Wash–Sagers Wash lenticular body), and upper Aberdeen Member turbiditic channel horizons are shown. Stops 1.1–1.4 are marked. The position of the Carmack Federal 1-14 (SWNE-14-21S-20E) drillhole is located ~2 km WNW of Stop 1.4. U.S. public land survey section boundaries and locations (i.e., section-township-range) are marked.

stones of the Desert, Grassy, and Sunnyside members underlie these sandstones and comprise the bulk of the cliff-forming strata. The combined thickness of the Castlegate Sandstone and Desert Member is 105–120 m, while the underlying Grassy and Sunnyside stratigraphic interval is 105–115 m thick (Fig. 2). These observations are consistent with well defined, memberscale thickness and sandstone content trends identified using nearby outcrop and subsurface data (Hettinger and Kirschbaum, 2002; Pattison, 2005a). The top of the Kenilworth Member occurs near the base of the cliff section and is marked by a slight increase in grain size (i.e., some siltstones and very fine-grained sandstones) relative to the overlying marine mudstones of the Sunnyside Member. This contact has been clearly identified in the subsurface and is expressed by a rightward deflection on the resistivity well logs and a leftward deflection on gamma ray and spontaneous potential well logs (Pattison, 2005a). The upper Kenilworth Member is dominated by marine mudstones and is ~50–55 m thick (Fig. 2). This interval is exposed in washes and low lying hills,

immediately south of the Book Cliffs. A 20–25-m cliff-forming succession marks the lower Kenilworth Member and is best defined north of Stop 1.4 (Fig. 3). An iron-rich siltstone occurs at the base of this cliff-forming succession, demarcating the Aberdeen-Kenilworth boundary (Figs. 4A and 4B). The lower Kenilworth cliff-forming package and the top Aberdeen iron-rich siltstone layer are the focus of Stop 1.2 (Fig. 3). The nearest cliff-forming package occurs a few hundred meters north of Stop 1.1 and consists of varying amounts of interbedded very fine-grained sandstones, siltstones, and mudstones (Figs. 4A, 4B, and 4C). These deposits are part of the upper Aberdeen Member and are separated from the top Aberdeen ironrich siltstones by a 25–30-m-thick package of marine mudstones. The cliff-forming package is nearly continuous across the Bootlegger Wash to Sagers Wash area, varying in height from 20 to 45 m (Fig. 3). At least two intervals of channel-rich deposits are observed (Figs. 2 and 4A), and these are equivalent to the X and Y intervals defined by Hampson et al. (1999). Both sandstone-rich (Stop 1.3) and mudstone-rich (Stop 1.4) channel-fill packages are

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Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior observed (Hampson, 2004). Good descriptions of the sedimentology and sedimentary architecture of Mancos Shale–encased channel-fill packages have been presented earlier (Chan et al., 1991; Cole and Young, 1991; Cole et al., 1997; Hampson et al., 1999; Hampson, 2004). Hampson et al. (1999) also provide paleocurrent data for X-Interval and Y-Interval channels and sheets, which indicates a dominant transport direction to the east or east-southeast (Table 1). Cumulative mi (km) 39.0

(62.8)

39.3

(63.2)

39.8

(64.0)

Description Return to vehicles and continue driving NW along the jeep track. Continue driving past a junction with a westerly-directed jeep track. Road swings to the north and winds through the wash. Park vehicles near the base of a Mancos Shale ridge/mound on the east side of the road. The base of this ridge is marked by an iron-stained horizon (top Aberdeen Member sequence boundary), while the upper half of the ridge is dominated by interbedded siltstone, mudstones, and very fine-grained sandstones (lower Kenilworth Member).

Stop 1.2—Bootlegger Wash–Sagers Wash: Thin-Bedded Turbidites and Iron-Rich Layer From top to bottom, the lower Kenilworth Member cliffforming package consists of (1) a 0.3-m-thick tightly ironcemented siltstone/mudstone; (2) 6.0 m of weathered mudstones with rare thin beds (a few cm) of siltstones and very fine-grained sandstones; (3) a 3.6-m-thick heterolithic package of interbedded very fine-grained sandstones, siltstones, and mudstones with an estimated sandstone content (net:gross ratio) of 20%; and (4) a 12.2-m-thick basal unit of weathered mudstones (Figs. 4A and 4B). Sandstone beds are sharp-based, 1–15-cm-thick, current-

Figure 4. Lower Kenilworth Member to upper Aberdeen Member stratigraphic interval in the Bootlegger Wash to Sagers Wash area. (A) Composite measured section from Stops 1.2 and 1.3 (Fig. 3); mst— mudstone; slst—siltstone; vf—very fine-grained sandstone; f—finegrained sandstone; m—medium-grained sandstone. Scale in meters. (B) Stop 1.2. Lower Kenilworth Member thin-bedded turbidites are concentrated in a 4-m-thick zone (above the scale bar) and are overlain by 6 m of weathered marine mudstones that are capped by an iron-rich siltstone layer (i.e., at the skyline). The top of the Aberdeen Member is observed near the base of the photo and is marked by a brownish-red iron-cemented siltstone layer (white arrows). (C) Stop 1.3. Sandstonerich turbiditic channel-fill deposit. Base is marked by a line and two white arrows. (D) Stop 1.4. Composite turbiditic channel-fill with a lower mudstone-rich package truncated by an overlying sandstonerich package. Key surfaces are highlighted by lines. People for scale at the base of the cliff.

485

rippled, and laterally continuous for tens to hundreds of meters. Planar-laminations are present at the base of some of the thicker sandstone beds. Groove marks occur at the base of some sandstone beds, while trace fossils are observed on upper bedding planes. Bioturbation is confined to the tops of the current-rippled sandstone beds and is sparse to moderately abundant. Most trace fossils are linear to slightly curved (2–10 cm long), narrow (0.2–0.5 cm) tracks or trails with chevron markings that resemble Scolicia or Gyrochorte, similar to those described elsewhere (Saunders and Pemberton, 1986; Frey and Howard, 1990; Pemberton et al., 2001). Rare Paleophycus and Thalassinoides trace fossils are also present. Paleocurrent data measured from the current-rippled sandstones indicates a mean transport direction of N101°E (n = 13) (Table 1). The base of the lower Kenilworth Member cliff-forming package is marked by an iron-rich siltstone/mudstone layer that is traceable throughout the Bootlegger Wash to Sagers Wash area and marks the contact between the Kenilworth and Aberdeen members (Fig. 4A). This layer is 0.1–0.3 m thick, with localized thickening up to 1.7 m. Thicker iron-rich siltstone layers appear to be concentrated as isolated pods or channels within this interval (Fig. 4B). These pods have concentrated layers of shell fragments, including ammonoids (Baculites) and brachiopods, as well as discontinuous lenses of very fine- to fine-grained sandstones. Cumulative mi (km) 39.8

(64.0)

40.3

(64.8)

40.8

(65.6)

Description Return to vehicles and turn back toward the south. Travel back through the wash. Turn right (W) onto a jeep trail. Drive carefully and slowly along this deeply rutted trail. Park vehicles at the base of a prominent Mancos Shale ridge that is capped by a sandstonerich channel-fill deposit. Jeep trail terminates at a man-made reservoir embankment. Climb and walk the ridge with caution: rattlesnakes have been previously encountered at this stop.

Stop 1.3—Bootlegger Wash–Sagers Wash: Sandstone-Rich Channel Package A sandstone-rich channel package outcrops along the south side of a west to east oriented, 200-m-long ridge (Fig. 3). The sandstone-rich channel package is ~6 m thick and is overlain by a 3–4-m-thick heterolithic package (10%–20% sandstone), which in turn is capped by a 0.3–0.8-m-thick, tightly cemented, iron-rich siltstone (Fig. 4A). The multistory channel package is sharp-based with up to 3 m of erosion (Fig. 4C). Internally, it comprises a stack of at least three nested channel-fill units, 1–3 m thick, with sharp and erosive basal contacts. Concaveup sandstone beds drape the margins of each channel-fill unit and these have low angle dips (2° to 7°). Onlap, truncation, and pinch-and-swell geometries are common. Some channel fills

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consist of fining-upward successions with the thickest sandstone beds near the base. Sigmoidal-shaped geometries are rare. Current ripple-laminated very fine- to fine-grained sandstone beds are the dominant facies within the sandstone-rich channel-fill package. Sandstones are well sorted. Combined-flow ripple-laminated sandstones are occasionally observed. Planar-laminated sandstone beds, some of which have gently curving lamination with or without small-scale (mm) truncation surfaces, are associated with the thicker sandstone beds. Normally graded beds consisting of a lower planar-laminated division and an upper ripplelaminated division are present. Sandstone beds are up to 1 m thick, with most in the 5–15 cm range. These beds are sharp-based, with rare sole marks, and have a concave-up geometry that drapes the channel margins. Convolute bedding is visible, especially in the steeper dipping portions of the inclined strata. Some beds show evidence of lateral accretion. Thinly bedded heterolithic facies consisting of interbedded mudstones, siltstones, and very fine-grained sandstones occurs between the sandstone beds. Carbonaceous-rich laminae, consisting of finely comminuted plant matter, and iron-rich laminae are abundant. Heterolithic strata are gently inclined, with concave-up and draping geometries. Truncation surfaces are visible. Bioturbation is mild to moderate and is mostly concentrated along the tops of the sandstone beds as tracks and trails. Recognizable trace fossils include Scolicia, Planolites, Paleophycus, Chondrites, Teichichnus, and Skolithos. Interpretations of Mancos Shale–encased channel-fill deposits are diverse and include turbiditic channels, shelf channels, prodelta to delta front channels, and tidally influenced, fluvial and estuarine channels deposited within incised valleys. The sandstone-rich channel package at Bootlegger Wash was previously interpreted as a valley-fill deposit (O’Byrne and Flint, 1993, 1995, 1996) or a tidally influenced fluvial channelfill deposit (Hampson et al., 1999) and was correlated to Grassy sequence boundary 2 (O’Byrne and Flint, 1993, 1995, 1996). Recent high-resolution outcrop and subsurface correlations indicate that this package is at least 125–150 m beneath the base of the Grassy Member, and is part of the upper Aberdeen Member (Fig. 2) (Pattison, 2005a, 2005d).

The overall package is ~15 m thick, with at least three nested mudstone-rich channel-fills, capped by two sandstone-rich channel-fills. Discrete channel-fills are sharp-based and 2–7 m thick. The sedimentary architecture of each channel-fill is characterized by concave-up geometry, beds draping channel margins, lateral accretion, onlap, and truncation surfaces (Fig. 4D). Individual channel-fills have a similar suite of facies compared to those at Stop 1.3, with the main difference being the higher proportion of mudstones and thin bedded heterolithic layers. Mudstone-rich channel packages are only revealed in outcrop localities where they are capped by a resistant rock type, such as a sandstone or iron-cemented siltstone bed. Most Mancos Shale–dominated intervals are intensely weathered and form low lying hills or mounds. Cliff-forming mudstone packages, like that displayed at Stop 1.4, are rare. It is extremely difficult to study the sedimentology and sedimentary architecture of these mudstone-dominated intervals in the absence of cliff-forming sections. Numerous traverses through desert floor washes in the Green River to Sagers Canyon region have revealed a great complexity of Mancos Shale deposits, which includes mudstone-rich channel-fill packages, erosional scours (i.e., mudstone on mudstone contacts), and concave-up depositional surfaces. This information is critical for developing robust interpretations of depositional processes, environments, and history of Mancos Shale–encased sedimentary bodies. Cumulative mi (km)

Walk ~1 mi west along the desert floor to a prominent cliff-forming Mancos Shale ridge that is dominated by mudstone-rich channelfill deposits. Climb the ridge along a Mancos Shale spur and examine the sedimentology, sedimentary architecture, and erosional contacts of this nested channel complex.

Walk back (E) to the vehicles parked at Stop 1.3. Retrace route back to Green River. Drive east on the heavily rutted jeep trail. 41.3 (66.5) Turn right (SE) onto the jeep track at the Tjunction. Continue driving SE-E along the jeep track. 42.5 (68.4) Turn right onto the Sagers Canyon dirt road. 43.8 (70.5) Enter Sagers Wash and pass underneath a railway bridge. 45.0 (72.4) Turn right and drive west on the old highway. 45.5 (73.2) Turn left (S) onto the deeply rutted “paved” road that leads to I-70. 46.0 (74.0) Turn right (W) on the I-70 entry ramp. Drive west on I-70. 74.0 (119.1) Exit I-70 at the Exit 162 ramp. 74.3 (119.5) Turn right (NW) onto the access road for Green River (Main Street). Travel into Green River. 75.7 (121.8) Turn right into the Comfort Inn parking lot.

Stop 1.4—Bootlegger Wash–Sagers Wash: Mudstone-Rich Channel Package

DAY 2 TRIP LOG: HATCH MESA AND GRAY–TUSHER CANYON

The multistory channel package at Stop 1.4 is dominated by mudstones and thin bedded heterolithic layers (Fig. 4D).

Begin your mileage log as you depart from the Comfort Inn parking lot.

Cumulative mi (km) 40.8

(65.6)

Description

40.8

Description

(65.6)

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487

Figure 5. Detailed field map of the Hatch Mesa to Floy Wash area (Fig. 1). The top of the Castlegate Sandstone and the Hatch Mesa succession (i.e., KPS2) outcrop belt are highlighted. The location of Stops 2.1–2.4 and the measured section through the Hatch Mesa succession (Fig. 6A) are shown. Public land survey boundaries (i.e., sections) are labeled.

Cumulative mi (km) 0.0

(0.0)

0.6 3.3

(1.0) (5.3)

10.0

(16.1)

10.4

(16.7)

Description Exit the Comfort Inn parking lot by turning left (E) onto Main Street. Travel east. Turn left onto the old highway. The old highway passes under a railway bridge. Turn left (N) onto a dirt track. Climb up a small hill. Park vehicles on the desert floor.

Stop 2.1—Hatch Mesa: Overview Hatch Mesa is located in the southern part of the Book Cliffs ~18 km east of Green River (Figs. 1 and 5). The Hatch Mesa

succession is 1–2 km southwest of Hatch Mesa and forms a gently curving, NW-SE–trending outcrop belt, ~7 km long (Fig. 5). This succession is completely encased by the Mancos Shale (Figs. 2 and 6). The structural dip is ~3° to 4° NNE along the southern face of Hatch Mesa. From top to bottom, the southern face of Hatch Mesa consists of fluvial sandstones of the Castlegate Sandstone, Desert Member coastal plain deposits, thickbedded shoreface sandstones of the Desert Member, thin-bedded shoreface sandstones and shelf mudstones of the Grassy Member, and marine mudstones and siltstones of the Sunnyside and Kenilworth members (Figs. 2 and 6B). The top of the Kenilworth Member is well defined near the base of the Hatch Mesa cliff section by (1) a cliff-forming package of silty mudstones, siltstones and very fine-grained sandstones at the top of the Kenilworth Member; and (2) a distinct color change from yellow-gray at

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Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior the top of the Kenilworth Member to dark gray at the base of the Sunnyside Member (Fig. 6B) (Pattison, 2004, 2005a). The top Kenilworth Member to top Hatch Mesa succession interval is exposed on the desert floor, southwest of Hatch Mesa (Figs. 5 and 6B). Thinly bedded, wave-rippled, very fine-grained sandstones of Kenilworth parasequences 3 and 4 (KPS3–KPS4) are exposed at ground level throughout this zone (Pattison 2005a, 2005b). High resolution outcrop correlations have established the time equivalency of Kenilworth parasequence 2 (KPS2) and the Hatch Mesa succession (Figs. 2 and 6A) (Pattison, 2004, 2005a, 2005b, 2005c). Cumulative mi (km) 10.4

(16.7)

11.0

(17.7)

Description Continue driving northward toward the Hatch Mesa succession. Park the vehicles near the railway tracks. Cross the tracks with care (this is an active railway line). Head north toward the Hatch Mesa succession and climb a relatively steep spur on the western side of the canyon.

Stop 2.2—Hatch Mesa: Proximal Hatch Mesa Succession The Hatch Mesa succession is ~15 m thick at Stop 2.2 and has a maximum sandstone content of 40% (Figs. 5 and 6). A prominent color change marks the base of this succession, which is caused by an increase in siltstone, carbonaceous matter, and iron-rich cement. Four bedsets, 2–6 m thick, comprise the Hatch Mesa succession (Fig. 6A). The lowermost bedset forms a well defined coarsening-upward succession that is capped by two thick sandstone beds (1–3 m thick). These sandstone beds can be correlated across the outcrop belt and serve as marker beds. Each marker bed consists of a stack of two or more discrete sandstone beds, with evidence of erosion and scouring. The upper three bedsets comprise weakly defined coarsening-upward successions that are weathered back from the cliff front and are covered in scree and vegetation (Fig. 6A). These units are difficult to

Figure 6. Lower Kenilworth Member at Hatch Mesa. (A) Measured section from the proximal (western) side of the Hatch Mesa succession (Fig. 5). The top of KPS2 and bedset boundaries are highlighted. Scale in meters. See Figure 4 for legend of symbols and abbreviations. (B) Southwestern face of Hatch Mesa. The Hatch Mesa succession (foreground) outcrops along the desert floor ~1 km south of Hatch Mesa (background). Rock packages abbreviations: C + D—Castlegate Sandstone and Desert Member; G—Grassy Member; S—Sunnyside Member; K—Kenilworth Member. (C) Low angle, isotropic to weakly anisotropic hummocky cross-stratified sandstones. Hummocks (h) and swales (s) are clearly visible. (D) Wave-modified turbidite bed within the Hatch Mesa succession. This Bouma-like Tbc bed has a combined flow (cf) ripple-laminated division resting on top of a planar-laminated (p) basal division. Pen for scale is ~15 cm long.

489

correlate with any degree of certainty across the study area. The sandstone beds of the Hatch Mesa succession have a sheet-like to lobate geometry and are flat lying with no clinoforms (Fig. 6B). Flame structures, load casts, pinch-and-swell geometry, compensatory stacking patterns, small-scale truncation surfaces, and onlap are common. Erosional scours, up to 3 m deep and a few tens of meters wide, are concentrated in the proximal (western) part of the Hatch Mesa succession. In general, the sandstone content decreases basinward. However, some sandstone beds thicken in a basinward direction prior to thinning. The background deposits consist of pinstriped mudstones and sandy siltstones, with a paucity of wave ripples. These fairweather deposits are more abundant than the high energy event beds. Sandstone beds have sharp bases with numerous sole marks and consist of well sorted very fine- to fine-grained sands. Paleocurrent measurements indicate a net transport direction of N72°E (Table 1). There is a great diversity of vertical stratification styles and sedimentary structures. Most beds are dominated by ripplelaminated and planar-laminated sandstones. Combined-flow ripples are most common, with lesser amounts of current ripples and wave ripples. Thick packages of climbing ripples are concentrated in some intervals. Planar-laminated sandstones can be gently curving, with or without small-scale truncation surfaces. Some massive or structureless sandstone divisions are noted near the base of thick sandstone beds. Hummocky cross stratified (HCS) sandstones are locally abundant in the western part of the Hatch Mesa succession. Many are low angle, weakly anisotropic to isotropic, similar to those described elsewhere (Fig. 6C) (Myrow and Southard, 1996; Myrow et al., 2002). Hummocky cross stratified sandstones are also more common in the uppermost bedsets. Many sandstone beds stack to form normally graded successions that pass from planar- to ripple-laminated sandstones (Fig. 6D). Other beds show multiple stacked planar and ripple-laminated sandstones, with relatively diffuse contacts between each division. Convolute bedding and sand volcanoes are also observed. Lag deposits consisting of shell fragments (e.g., Baculites), ironstone pebbles, wood fragments, and coal clasts are rarely present at the base of the thickest sandstone beds, especially in the western (i.e., proximal) part of the outcrop belt. A low diversity and low abundance trace fossil suite, and locally abundant carbonaceous matter is present in all facies. There are a wide variety of interpretations for the lower Kenilworth Member isolated sandstone bodies. For example, the Hatch Mesa succession has been interpreted as a detached lowstand shoreface (Hampson et al., 1999, 2001), a delta front and distributary mouth bar (Taylor and Lovell, 1991, 1995), a shoreface attached delta lobe (Newman, 1985), a prodelta plume (Patterson, 1983; Swift et al., 1987; Cole and Young, 1991; Cole et al., 1997), a shelf deposit (Chan et al., 1991), a bottom-set position on a ramplike clinoform (Stevens and Chaiwongsaen, 2003; Stevens et al., 2005), and a storm-influenced prodelta turbidite complex (Pattison, 2004, 2005a, 2005b, 2005c; Pattison and Hoffman, 2005). The background deposits in the distal lower Kenilworth Member consist of non- to mildly bioturbated, pinstriped

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mudstones and sandy siltstones, with a paucity of wave ripples, suggesting deposition below fair weather wave base. In contrast, the sandstone beds represent high energy event beds that punctuate the quiet-water, background conditions. The sandstone facies have a great diversity of vertical stratification styles and sedimentary structures and are dominated by turbidite-like and stormgenerated event beds. Thick beds are Bouma-like Tbc or Tabc turbidites or HCS sandstones, while thin beds are predominantly Bouma-like Tc or Tbc turbidites or wave-rippled sandstones. HCS and wave ripple-laminated sandstones indicate deposition above storm wave base. The abundance of combined-flow ripple-laminated sandstones in the upper parts of the turbidite beds, coupled with lesser amounts of wave ripple-laminated sandstones and gently curving planar-laminated sandstones, suggest that these turbidites are wave-modified (Myrow and Southard, 1991, 1996; Myrow et al., 2002). The gently curving planar-laminated sandstone beds are very similar to the combined flow plane bed described by Arnott and Southard (1990). Stacked Bouma-like Tbc beds with relatively diffuse contacts between the beds suggest deposition from waning to waxing to waning hyperpycnal flows. Additional evidence supporting hyperpycnal flows is the locally abundant carbonaceous matter and the low diversity and low abundance trace fossil suite. Convolute bedding, sand volcanoes, flame structures, and load casts imply rapid deposition on top of water-saturated muds. The pinch and swell sand body geometry, compensatory stacking, and lack of clinoforms support deposition in a bottom-set position of a ramp-like clinoform (Stevens and Chaiwongsaen, 2003; Stevens et al., 2005). Cumulative mi (km) 11.0

(17.7)

12.0 13.9

(19.3) (22.4)

14.2

(22.8)

14.8 15.0

(23.8) (24.1)

Description Return to vehicles and drive south on the dirt track toward the old highway. Turn left (E) onto the old highway. Turn left (N) onto the Floy Canyon road (note the sign). Stop before crossing the railway tracks. This is a non-signaled “level crossing” that is very active with railway traffic. Proceed with caution. Drive through Floy Wash. Pull vehicles over and park. Walk a short distance west of the Floy Wash road to the low lying outcrops of the Hatch Mesa succession.

Stop 2.3—Hatch Mesa–Floy Wash: Distal Hatch Mesa Succession This stop is located ~2.6 km down–depositional dip of Stop 2.2 (Fig. 5). The Hatch Mesa succession is ~10 m thick and has a sandstone content of 20%. The two marker beds capping the lowermost bedset have dramatically thinned to 0.1–0.4 m. These marker beds are no longer the thickest sandstone beds in the Hatch Mesa succession, as the thickest beds are present

in the uppermost bedsets. A similar suite of facies and vertical stratification styles is present in the distal Hatch Mesa succession as compared to the proximal areas. One significant difference is the greater proportion of ripple-laminated sandstones at Stop 2.3 compared to Stop 2.2. Combined-flow and current ripple-laminated sandstones dominate the lower half of the succession. The upper half of the succession has a mixture of sandstone beds including low-angle HCS sandstones, gently curving planar laminated sandstones, and planar- to ripple-laminated sandstone beds. A coarse-grained package outcrops sporadically on the desert floor near the base of the cliff section. This stratigraphic interval is the focus of Stop 2.4. Cumulative mi (km) 15.0

(24.1)

15.1

(24.3)

15.3

(24.6)

Description Return to the vehicles and do a U-turn. Travel south on the Floy Wash road. Turn left (E) onto a dirt track and follow it to its termination. Park vehicles and walk a short distance southeast.

Stop 2.4—Hatch Mesa–Floy Wash: Coarse-Grained Sandstone and Lag Deposit Stop 2.4 is located ~0.5 km down–depositional dip of Stop 2.3 (Fig. 5). The Hatch Mesa succession maintains a thickness of ~10 m, although there is less sandstone compared to the previous stop. The primary focus of Stop 2.4 is to examine the coarse-grained sandstone and lag deposit that occurs along the Aberdeen-Kenilworth boundary (Fig. 2). A secondary focus of this stop is to examine the oolitic ironstone deposit a few meters beneath the coarse-grained sandstone. Both rock types have been described in earlier studies (Chan et al., 1991; Chan, 1992; Taylor and Lovell, 1995; Cole et al., 1997; Hampson et al., 1999; Hampson, 2004). There are numerous isolated pods of coarse-grained sandstones in the Floy Wash area; these vary in thickness from 0.5 to 3.5 m. These pods have a limited areal extent and are typically a few meters to a few tens of meters wide and long. Three-dimensional sand body geometries include laterally accreted bar-like forms, channels, and ribbon-like bodies. A ribbon-like body at Stop 2.4 is oriented N124°E and terminates at a lobate-like body. Many of the coarse-grained sandstone bodies are mounded with convex-up geometries and lateral accretion surfaces with 2° to 10° dips. Most coarse-grained sandstone bodies have relatively sharp and erosive basal contacts, with truncation surfaces, sigmoidal bedding, pinch-and-swell geometry, and both convex-up and concave-up bedding planes. Some coarsening-upward successions are observed. Most of these bodies have a high proportion of sandstone, ranging from 60% to 90%. Fine- to mediumgrained sandstones dominate the lower half of each body, while medium- to very coarse-grained sandstones are common in the upper half. The sandstones have a salt-and-pepper texture and

Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior are moderately to poorly sorted. Angular to subangular mudstone clasts or chips are common (0.1–2.3 cm diameter). Granules and pebbles are scattered throughout the upper 0.5 m, with an increased concentration in the upper 0.1 m. Mudstone clasts, bone fragments, shell debris, and fish teeth are concentrated in the uppermost 0.1 m, forming a well-defined lag deposit. Wood fragments, coal fragments, and finely comminuted plant matter are identified in trace amounts throughout the coarse-grained bodies. Recognizable fossil fragments include Baculites and Inoceramus. Brachiopod shells, bivalve shells, and fish teeth are also present. The coarse-grained sandstone packages are mildly to moderately bioturbated, with the exception of the uppermost lag deposit, which is moderately to thoroughly bioturbated. Trace fossils include Paleophycus, Ophiomorpha, Thalassinoides, and Skolithos. Low-angle cross bedding, trough cross bedding, gently curving planar laminations, and current ripple laminations are the main sedimentary structures. Combined-flow and wave ripplelaminated sandstones are rare. Convolute bedding and normally graded beds are also observed. Paleocurrent data indicates a mean transport direction of N141°E (n = 33) (Table 1). These measurements are consistent with previously published paleocurrent data that indicated a mean transport direction of N159°E (n = 14) (Table 1) (Chan et al., 1991). An oolitic ironstone deposit is observed ~5 m below the top of the coarse-grained sandstone and lag deposit at Floy Wash (Chan, 1992). The oolitic ironstone caps low-lying mounds or hills and is up to 2 m thick. This deposit is ironrich, tightly cemented, and becomes coarser-grained upward. The upper 0.5 m is characterized by medium to coarse-grained oolites. Individual beds are 2–7 cm thick and are planar- or ripple-laminated. The oolitic sandstone beds are interbedded with iron-rich siltstones. Two conflicting interpretations have been proposed for the origin of the coarse-grained sandstone interval at the AberdeenKenilworth boundary. One interpretation links these deposits to sequence boundary development and subsequent transgressive reworking (Taylor and Lovell, 1991, 1995; Hampson et al., 1999, 2001), while the other interpretation proposes submarine erosion and transgressive winnowing in an inner shelf environment (Balsley, 1980; Newman, 1985; Swift et al., 1987; Chan et al., 1991; Chan, 1992; Cole et al., 1997). The balance of the evidence presented herein supports the former interpretation. The coarse-grained deposits are interpreted as falling stage- and lowstand-derived sandstone bodies that mantle a sequence boundary. Coarse-grained sediments would have been brought onto the exposed shelf during falling stage and lowstand through a network of incised fluvial channels. The thin (0.5–3.5 m) and patchy occurrence of the coarse-grained bodies reveals that the depth of incision was relatively shallow. The valley systems were likely similar to the “unincised valleys” described by Posamentier (2001). The coarse-grained sandstone bodies were later reworked and possibly remolded during subsequent transgressive ravinement. Coarse-grained deposits are restricted to a few scattered localities. Elsewhere, this transgressively modified sequence

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boundary is marked by iron-rich siltstone and mudstone beds (Stop 1.2), with ammonoid (Baculites) and brachiopod shell fragments and discontinuous lenses of very fine- to fine-grained sandstones. These represent the interfluves of the transgressively modified sequence boundary. In summary, the top Aberdeen sequence boundary is demarcated by a variety of deposits including marine mudstone-encased iron-rich siltstones, coarse-grained sandstones, and coarse-grained lag deposits (i.e., pebbles, fish teeth, shell fragments, bone fragments, and mudstone clasts). Paleocurrent data indicate a N145°E transport direction and, therefore, the lowstand shoreline was likely southeast of the study area (Table 1). Cumulative mi (km) 15.3

(24.6)

15.5 15.6 16.2

(24.9) (25.1) (26.1)

16.5 16.8

(26.5) (27.0)

27.8 28.1

(44.7) (45.2)

29.3

(47.1)

35.9

(57.8)

36.9

(59.4)

37.2

(59.9)

Description Return to the vehicles. Drive west toward the Floy Wash road. Turn left (S) onto the Floy Wash road. Drive through Floy Wash. Cross the railway tracks with care. Look both ways before crossing. Turn left (E) onto the old highway. Turn right onto the I-70 (W) entrance ramp (Ranch Exit 173). Travel west on I-70 toward Green River. Exit I-70 via the Exit 162 ramp. Turn right (NW) onto the access road for Green River (Main Street). Travel into Green River. Turn right onto a paved road and drive north. This road runs parallel to the Green River and is located along the eastern banks of the Green River. Drop down a steep hill. The road “bottomsout” in Tusher Wash. Continue driving north on this road. Turn right onto a dirt track that is signed for Tusher Canyon; follow the track east and south. Park the vehicles and climb the alluviumveneered Mancos Shale hill to the south.

Stop 2.5—Gray Canyon–Tusher Canyon: Overview This area is located on the east side of the Green River, ~12 km north of the town of Green River (Figs. 1 and 7). Stop 2.5 is ~1 km southwest of the Book Cliffs and is situated at the mouth of Tusher Canyon (Fig. 7). The Book Cliffs are ~300 m high, with the uppermost cliff-forming sandstones comprising the Castlegate Sandstone and Desert Member (Fig. 2). The top of the Desert Member shoreface package is marked by a prominent “whitecap” sandstone interval (Fig. 8A). Whitecap intervals have been identified throughout the Blackhawk Formation in east-central Utah and are formed by the leaching

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of detrital dolomite and feldspar clasts by acidic waters sourced from the overlying coals (Spieker, 1931; Young, 1955; Balsley, 1980; Taylor et al., 2000). The combined thickness of the Castlegate Sandstone and Desert Member is 75 m. The Grassy Member is the next significant cliff-forming sandstone package and is 65 m thick. Grassy parasequence 2 (GPS2) has a whitecap at its top (Fig. 8A). This is overlain by a thin screecovered package representing GPS3 and GPS4. The Sunnyside Member is dominated by marine mudstones and siltstones, with the uppermost 15–20 m represented by a cliff-forming package of interbedded siltstones and very fine-grained sandstones. This member is 80 m thick. All four Kenilworth parasequences are visible along the east side of the Green River near the base of the Book Cliffs. The upper part of KPS4 is identified by a color change in the scree-covered and weathered outcrop that passes from yellowgray mudstones and siltstones in the uppermost Kenilworth Member to dark gray mudstones in the lowermost Sunnyside Member. The middle and lower parts of KPS4 and the upper parts of KPS3 and KPS2 are identified as cliff-forming packages. Kenilworth parasequence 1 (KPS1) is present near the base of the KPS2 cliff section and is only visible along a narrow stretch of outcrop. The Kenilworth Member is ~100 m thick in this region, with the top Kenilworth Member to top KPS2 stratigraphic interval encompassing 65 m. Lobate inner shelf sandstone bodies are present in the upper parts of KPS2 and KPS1. The Aberdeen-Kenilworth contact is recognized along an ironrich siltstone layer that is present near the base of the Book Cliffs. This layer sporadically outcrops in the Gray Canyon to Tusher Canyon area and is correlative with the coarse-grained sandstones examined at Floy Wash (Stop 2.4). Channel-fill deposits are present at two distinct levels in the upper Aberdeen Member, and these are exposed on the desert floor near the base of the Book Cliffs (Figs. 8B, 8C, and 8D). Cumulative mi (km) 37.2

(59.9)

Description Walk ~0.4 mi northeast to the low-lying outcrops on the desert floor.

Figure 7. Detailed field map of the Gray Canyon to Tusher Canyon area (Fig. 1), characterized by turbiditic-rich channel-fill deposits along two horizons (Levels 1 and 2) in the upper Aberdeen Member. The top of Kenilworth parasequence 2 (KPS2) is also shown. Stops 2.5 and 2.6 are located in the SW corner. Public land survey boundaries (i.e., sections) are labeled.

Stop 2.6—Gray Canyon–Tusher Canyon: Upper Aberdeen Member Channel-Fill Deposits Upper Aberdeen Member channel-fill deposits are concentrated along two horizons at Stop 2.6 (i.e., Levels 1 and 2). Level 1 channel-fill deposits are 6–12 m thick and have a higher sandstone content compared to Level 2 channel-fills, which are 3–7 m thick (Fig. 8B). These deposits are concentrated along the same stratigraphic levels as the channel-fill deposits in the Bootlegger Wash to Sagers Wash area (Stops 1.1–1.4) and the Gunnison Butte to NW Willow Bend area (Stops 3.1 and 3.2). Channels are sharp-based and consist of a multistory package of nested channel-fill units. Well-defined truncation surfaces separate individual channel-fill units, which consist

of mudstone-, heterolithic-, or sandstone-rich fills. These units are dominated by concave-up, sigmoidal-shaped sand body geometries, which exhibit onlap and downlap (Fig. 8B). Beds drape the channel margins. Sandstone beds are sharp-based with sole marks, and consist of well sorted, very fine- to finegrained, ripple-laminated, planar-laminated, or low-angle cross bedded sandstones. Paleocurrent data indicate a transport direction to the east-southeast (Table 1) (Chan et al., 1991). Some sandstone beds have moderately bioturbated tops, with Scolicia and Paleophycus being the most common trace fossils. Carbonaceous-rich laminations consisting of finely comminuted plant material are locally abundant.

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Figure 8. (A) Photo-panorama of the Book Cliffs on the east side of the Green River between the mouth of Gray Canyon (left) and the NW entrance to Tusher Canyon (right). Kenilworth parasequences 1–4 (KPS1 to KPS4) are labeled. Level 2 channels (upper Aberdeen Member) are highlighted in the lower right. Rock package abbreviations as per Figure 6. (B) Photo-panorama of the upper Aberdeen to lower Kenilworth stratigraphic interval. View is toward the NE. Turbiditic-rich channels occur along two horizons in the upper Aberdeen Member (i.e., Levels 1 and 2). The KPS2 cliff-line is in the background. (C) Stop 2.6. Level 1 channel complex. Channel base is marked by a black line. (D) Level 1 channel complex.

S.A.J. Pattison

494 Cumulative mi (km) 37.2

(59.9)

37.5

(60.3)

38.5 45.1 45.3

(61.9) (72.6) (72.9)

Description Return to the vehicles and retrace the route back to Green River. Drive north and west along the Tusher Canyon access road toward the paved road. Turn left (S) onto the paved road, heading to Green River. Tusher Wash. Turn right (W) onto Main Street. Turn right into the Comfort Inn parking lot.

DAY 3 TRIP LOG: GUNNISON BUTTE AND PRICE RIVER CANYON Begin your mileage log as you depart from the Comfort Inn parking lot. Cumulative mi (km) 0.0

(0.0)

0.2 1.3

(0.3) (2.1)

7.2

(11.6)

7.5

(12.1)

Description Exit the Comfort Inn parking lot by turning right (W) onto Main Street. Travel west. Bridge over the Green River. Turn right (N) onto Long Street. Follow this road north out of town. This road has a number of right-angle bends. Long Street goes from a paved road to a gravel road. Continue driving north. Pumping station. The Green River is visible on the right side (E) of the road. Road twists and turns.

8.1

(13.0)

9.0

(14.5)

9.5

(15.3)

9.8

(15.8)

Drive past the Blue Castle access road (dirt track to the W). Road turns from north to east and follows the banks of the Green River. Abrupt right angle turn to the north away from the Green River. Continue to follow the road north through the Gunnison Butte ranch property. Turn left onto a westward-trending access road on the northern edge of an irrigated field (get permission from the landowners). Drive slowly and carefully westward (i.e., avoid damaging irrigation pipes and equipment). Park the vehicles and climb over the barbed wire fence toward the north. Walk to the top of the small hill.

Stop 3.1—Gunnison Butte–NW Willow Bend: Overview Gunnison Butte is a prominent landform situated along the western mouth of Gray Canyon, which marks the passage of the Green River onto the broad desert landscape south of the Book Cliffs (Figs. 1 and 9). Willow Bend is a densely vegetated part of the Green River valley directly south of Gunnison Butte (Fig. 9). The east Gunnison Butte spire comprises cliff-forming sandstones of the Castlegate Sandstone and Desert Member (75 m thick). Directly west of the spire are the flat-topped western and central parts of Gunnison Butte, which is capped by the top of the Grassy Member. Both Grassy parasequences (i.e., GPS1 and GPS2) have a high proportion of sandstone at this locality. The approximate thicknesses of the Grassy, Sunnyside, and Kenilworth members are 50 m, 80 m, and 100 m, respectively (Fig. 2).

Figure 9. Detailed field map of the Middle Mountain to Gunnison Butte area showing the location of Stops 3.1–3.3. The base of the Book Cliffs is marked by the KPS2 cliff-line. Level 1 and Level 2 turbiditic channel-fill deposits outcrop on the desert floor a few hundred meters south of the main cliff-line. These upper Aberdeen Member deposits are concentrated in a relatively small area.

Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior Kenilworth parasequences 2, 3, and 4 make up the lower three cliff-forming benches in the Middle Mountain to Gunnison Butte corridor (Figs. 10 and 11). All three parasequences have interbedded sandstones and siltstones near the top, which progressively thin and become muddier toward the east. The top of the Kenilworth Member (i.e., top KPS4) is present 8 m above the uppermost cliff-forming bench and is demarcated by a color change in the scree-covered mudstones. This color change is well defined between Tusher Canyon and Hatch Mesa, which allows for the confident correlation of the top Kenilworth Member contact across this area (Pattison, 2004, 2005a, 2005b). A retrogradational bedset stacking pattern characterizes the uppermost Kenilworth Member in this region (Pattison, 1995; Hampson, 2000). KPS2 has an anomalously high sandstone content in the Middle Mountain to Gunnison Butte region, and is dominated by turbidite-like event beds and HCS sandstones (Fig. 10). Compared to the Hatch Mesa succession (i.e., Stops 2.1–2.3), the Middle Mountain–Gunnison Butte lenticular body has a similar suite of facies and sedimentary architecture, and is also at the same stratigraphic level (KPS2) (Pattison, 2005a, 2005b). This lenticular body will be the focus of Stop 3.3. KPS1 is near the base of the Book Cliffs in the Gunnison Butte region. The top of this parasequence is marked by interbedded siltstones and very fine-grained sandstones, which are highlighted by a color change in outcrop. The Aberdeen-Kenilworth contact is at the base of the Book Cliffs and is marked by a variety of deposits in this area including (1) an iron-rich siltstone bed, (2) fine-grained channel-fill deposits, or (3) a coarse-grained sandstone bed capped by a lag deposit (Fig. 10). The best example of the latter is beneath the aptly named “Toad Stool” (Lee Krystinik, 2004, personal commun.) landmark in the western part of the study area (Figs. 9, 10, and 11A). These deposits reveal a transport direction of N134°E (n = 6) (Table 1). The coarse-grained sandstone bed with upper lag deposit will not be visited during this field trip but can be viewed with binoculars from Stop 3.2. Channel-fill deposits are present at two distinct levels (Levels 1 and 2) in the upper Aberdeen Member, forming a series of low lying ridges and hills along the desert floor NW of Willow Bend (Figs. 9 and 11). These channel-fill deposits are time equivalent to those examined on the east side of the Green River at Stop 2.6 (Figs. 2 and 10). Cumulative mi (km) 9.8

(15.8)

Description Walk ~0.5 mi to the channel-fill outcrops.

Stop 3.2—Gunnison Butte–NW of Willow Bend: Upper Aberdeen Member Channel-Fill Deposits Upper Aberdeen Member channel-fill deposits are concentrated in a 0.4 km2 area NW of Willow Bend and are observed at two distinct horizons (Fig. 9). Level 2 is 20 m below the top of the Aberdeen Member, while Level 1 is 15 m lower (Fig. 10). The best developed channel-fill deposits are present along Level

495

2, and these are 2–7 m thick (Figs. 10 and 11B). Well developed sandstone-rich channel-fill units are present along the Level 2 horizon (Fig. 11B). In contrast, Level 1 channel-fill deposits are 1–5 m thick and are dominated by mudstone- and heterolithicrich channel-fill units. Each multistory channel-fill package is a few tens of meters wide and consists of 2–5 nested and stacked channel-fill units, with sharp and erosional contacts between each unit. Most units are dominated by concave-up beds that drape the channel margins. Lateral accretion surfaces are common, with dips ranging from 2° to 11°. Individual sandstone beds show pinch-and-swell geometries, onlap, and downlap. The dominant facies within the channel-fill packages are pinstriped mudstones, siltstones, and very fine- to fine-grained sandstones. Finer grained facies have an abundance of finely comminuted plant material and show a low degree of bioturbation. Thin, weakly rippled sandstone beds and normally graded siltstone layers are common within the pinstriped mudstones. Some siltstone and mudstone beds are tightly cemented, with iron-rich nodules concentrated along the base of some of channels. The sandstone beds are sharp-based, 1–50 cm thick, are dominated by ripple-laminated and planar-laminated sandstones, and are well sorted. Both current and combined flow ripples are observed. Starved current ripples and climbing ripple sets are locally abundant. Some planar-laminated sandstone beds are gently curving or undulating, with long wavelengths (0.5–1.5 m) and low amplitudes (a few cm), with or without low angle truncation surfaces (Fig. 11C). Planar-laminated sandstone beds are often overlain by ripple-laminated sandstone beds. Stacked sets of planar- and ripple-laminated sandstones, with relatively diffuse contacts between each division, are also observed, with a maximum 12 divisions in one stack (Fig. 11D). Minor amounts of wave ripples, low angle HCS sandstones, and convolute bedding are present. Current ripple and channel axes paleocurrent data reveals a N114°E (n = 107) transport direction for Level 2 channels and a N122°E (n = 58) transport direction for Level 1 channels (Table 1; Fig. 12B). Organic-rich laminations, consisting of finely comminuted plant material, are observed in all sandstone facies. Moderate bioturbation is observed on some bedding planes, with Scolicia, Paleophycus, and Thalassinoides. Trace fossils rarely cut across laminations or beds. Low concentrations of Inoceramus and Baculites fossil fragments are present in most channel-fill units. One Level 2 channel-fill package has a 0.5-m-thick transgressive lag deposit, with an unusually dense concentration of Inoceramus and Baculites fragments, some up to 20 cm long. The sharp-based planar- to ripple-laminated sandstone beds with sole marks are interpreted as Bouma-like Tbc beds (Pattison et al., 2005). Long-wavelength, low-amplitude, gently curving planar-laminated sandstone beds with truncation surfaces are interpreted as low angle HCS sandstones. The presence of the low angle HCS sandstones, combined-flow ripple-laminated sandstones, and rare wave ripple-laminated sandstones indicates deposition above storm wave base (Pattison et al., 2005). The abundance of finely comminuted plant material, low diversity

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Figure 10. High resolution correlation of the Kenilworth Member and upper Aberdeen Member in the Middle Mountain to Tusher Canyon area. Datum is the top of the Aberdeen Member, which is marked by a coarse-grained transgressive lag deposit. Kenilworth parasequences (KPS1 to KPS4) and the upper Aberdeen turbiditic channel levels (L1 and L2) are labeled. The locations of the four measured sections are highlighted on the inset map. Line of projection is west to east. Distances between each section given in km. Vertical scale in meters. See Figure 4 for a legend of symbols and abbreviations.

Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior trace fossil assemblage, and a mild degree of bioturbation indicate freshwater-terrestrial input. Stacked sets of Bouma-like beds are interpreted as hyperpycnal-flow–derived turbidites or hyperpycnites. Concentrations of Inoceramus and Baculites shell fragments in the upper parts of channel-fill packages are interpreted as transgressive lag deposits, implying that wave base impeded on the tops of some channel-fill deposits. The variable sandstone content, multistory sedimentary architecture, and sharp and erosive contacts between nested channel-fills indicate a wide range in energy conditions on the inner shelf. Many channel-fill units are fine-grained and are dominated by mudstones and siltstones. These flows would not have been capable of cutting large-scale channel incisions. Therefore, it is highly probable that the cutting and filling of these channels was achieved during a series of discrete events separated by time. In other words, high-energy flows that led to channel incision on the inner shelf would also have transported their sediment load much further basinward, leading to sediment bypass. These channels would be subsequently infilled by low-energy flows that were not only too weak to incise into the inner shelf, but were also incapable of bypassing sediments to the middle and outer shelf. Further support for low-energy channel infilling events is provided by the predominantly concave-up sandstone bed architecture that conformably drapes the channel margins. These sandstone beds mimic the channel morphology and are interpreted as a lowenergy “passive” infill deposit. Cumulative mi (km) 9.8

(15.8)

10.1

(16.3)

10.3

(16.6)

10.5

(16.9)

11.0

(17.7)

Description Return to the vehicles and do a U-turn. Drive east on the access road. Merge with Long Street and continue to drive east. Turn left (N) at the T-junction; drive toward Gunnison Butte on the west side of the corral. Turn right (E) and drive through the corral. You will soon drive past an orchard on the north side of the dirt track. Continue driving until the road becomes impassable. Park the vehicles. Walk a short distance north along the dirt track, which skirts the eastern boundary of Gunnison Butte. The rocks exposed at road level are KPS2 (lower Kenilworth Member).

Stop 3.3—Eastern Gunnison Butte: Lower Kenilworth Member Inner Shelf Deposits Both KPS2 and KPS1 have an anomalous concentration of sandstone, which is up to 40% in some parts of the Middle Mountain to Tusher Canyon region (Figs. 2 and 10) (Pattison, 1995, 2005a). This is significantly greater than proximal areas to the west, where the sandstone content rarely exceeds 10%. Lobate geometries are particularly well defined along the KPS2

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horizon, especially at Gunnison Butte (Fig. 11E) and along the east side of the Green River (Fig. 8A). The purpose of Stop 3.3 is to examine the sedimentology of the Middle Mountain–Gunnison Butte lenticular body on the east side of Gunnison Butte (Fig. 9). The Middle Mountain–Gunnison Butte lenticular body has a wide variety of sandstone beds, including Bouma-like Tbc, Tabc, or Tc beds (Fig. 11F), combined-flow ripples, current ripples, HCS sandstones, gently curving planar laminations, wave ripples, large-scale low-angle symmetrical to weakly asymmetrical ripple-laminated sandstones, massive sandstones, convolute bedding, and thick sets of climbing ripple laminations. Most of these sandstone beds are sharp-based with well defined sole marks and consist of well-sorted, very fine- to finegrained sands. Paleocurrent data from sole marks and current or combined-flow ripples show transport toward the east, along a N72°E trend (Table 1; Fig. 12D). Some Bouma-like beds stack in sets of two or more beds. A few have ripple-laminated sandstones at the base that pass upward into planar-laminated sandstones. These deposits are characterized by a low to moderate degree of bioturbation, with recognizable trace fossils including Scolicia, Chondrites, Helminthopsis, Ophiomorpha, Paleophycus, and Planolites. Some intervals are non-bioturbated. Finely comminuted plant material is common throughout all facies. Cumulative mi (km) 11.0

(17.7)

11.5 11.7

(18.5) (18.8)

13.3

(21.4)

13.9 14.2

(22.4) (22.8)

20.1

(32.3)

21.6

(34.8)

23.8

(38.3)

49.7 49.9 50.0

(80.0) (80.3) (80.5)

54.4

(87.5)

Description Return to the vehicles and retrace the route back to Green River. Turn left (S). Turn right (W). Follow Long Street back to Green River. Drive past the Blue Castle access road (dirt track to the W). Continue driving south. Pumping station. Long Street changes from a gravel road to a paved road. Continue driving south. Turn right (W) onto Main Street. Drive west through Green River. Turn right onto the on-ramp for I-70 (W). This is Exit 158. Travel west on I-70. Turn right off I-70 (Exit 156) and drive northwest onto Hwy 6/191. Continue driving north on Hwy 6/191. Price River. Gas station on west side of road at Woodside. Turn right (E) onto a dirt road. This is the Price River Canyon access road. Continue driving east on this road. Drive with caution as this road has many blind corners and hills. After passing a Mancos Shale hill on the left side of the dirt road, turn left and park the vehicles in an open area.

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Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior Stop 3.4—Price River Canyon: Overview The final field trip area is located along the north bank of the Price River, ~6 km east of Woodside (Fig. 1). The Price River flows eastward into the Book Cliffs and joins the Green River some 18 km farther downstream. The Book Cliffs are ~300 m high in the Price River Canyon area. The Castlegate Sandstone is the uppermost cliff-forming package and is up to 90 m thick (Fig. 2). This interval rests on top of a 25-m-thick package of scree-covered coastal plain deposits of the Desert and upper Grassy members. A 90-m-thick stratigraphic interval with four well-developed whitecap intervals is present beneath the coastal plain deposits. These whitecap intervals are at the top of Grassy parasequence 1 (GPS1) and at the top of all three Sunnyside parasequences. The top of the Kenilworth Member (i.e., KPS4) is marked by a 25-mthick package of upper shoreface sandstones (Fig. 2). Two other cliff-forming units are observed below the thick KPS4 sandstone package; these represent KPS3 and KPS2. KPS2 is the focus of Stop 3.5, while KPS4 is focus of Stop 3.6 (Fig. 1). Cumulative mi (km) 54.4

(87.5)

Description Begin hiking north into the side canyon. Pack water and watch out for rattlesnakes. A moderate hike and climb will bring you to the base of the first cliff-forming unit, which is Mancos Shale–dominated and is within the lowermost part of the Kenilworth Member. Continue hiking along the western wall of the side canyon

Figure 11. Middle Mountain to Gunnison Butte region. (A) Location of the Toad Stool measured section shown in Figure 10. A coarse-grained sandstone package with an upper transgressive lag deposit caps the top of the Aberdeen Member. Kenilworth parasequences (KPS) 2–4 are labeled. (B) Level 2 channel complex consisting of at least three channel-fill packages: a lowermost mudstone-rich unit cut by two heterolithic units. Location is ~100 m northwest of Stop 3.2 along the same ridge-line. (C) Level 2 channel-fill deposit along the eastern side of the outcrop belt, NW of Willow Bend (Fig. 9). Channel-fill is characterized by weakly asymmetrical large-scale ripple-laminated fine-grained sandstones with low angle dips and truncation surfaces. This facies is interbedded with silty mudstones and weakly rippled very fine-grained sandstones. (D) Level 2 channel-fill deposit from the central part of the outcrop belt. Interbedded current ripple-laminated (c) and planarlaminated (p) sandstones. At least nine separate divisions are recognized within a 30-cm-thick bed. Many of the planar-laminated divisions show gently curving laminations. Pen for scale (15 cm long). (E) The Middle Mountain–Gunnison Butte lenticular body showing a subtle decrease in sandstone content and a gentle stratigraphic dip (≈2°) toward the north (left). Level 2 deposits in the upper Aberdeen Member are located in the foreground at the base of the photograph. (F) Bouma-like Tabc bed from KPS2 along the eastern side of Gunnison Butte near Stop 3.3 (Fig. 9). The massive (m) division is the thickest and is overlain by thin planarlaminated (p) and combined-flow ripple-laminated (cf) divisions. Note the symmetry of some combined-flow ripple crests (arrows).

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until you reach an undercut part of the cliff front. Be careful near the undercut cliff. Blocks have recently been dislodged. The interbedded sandstones, siltstones, and mudstones are part of KPS2 (lower Kenilworth Member). Stop 3.5—Price River Canyon: Lower Kenilworth Member Inner Shelf Deposits The lack of time equivalent outcrop directly west of Price River Canyon makes it impossible to determine if the lower Kenilworth Member inner shelf deposits are shoreface attached or detached. These inner shelf deposits have a sedimentology similar to the KPS2 deposits at Gunnison Butte (Stop 3.3). KPS2 is subdivided into two coarsening-upward bedsets, 10–15 m thick, with 30%–50% sandstone. The sandstone beds are sharpbased, with sole marks, and are dominated by very fine- to finegrained HCS sandstones, Bouma-like Tbc beds, gently curving planar laminations, and combined-flow ripple laminations. Load casts, flame structures, and convolute bedding are associated with the meter-scale sandstone bed that caps the lower bedset. Organic-rich laminations, consisting of finely comminuted plant debris, and non-bioturbated intervals are common. Cumulative mi (km) 54.4

(87.5)

Description Continue hiking and climbing up the canyon. Compare and contrast the thick bedded sandstones in the upper part of the Kenilworth Member with the thinner bedded sandstones in the lower Kenilworth Member.

Stop 3.6—Price River Canyon: Upper Kenilworth Member Shoreface Deposits The upper Kenilworth Member comprises two parasequences: KPS3 and KPS4. These intervals show a strong wave- and storm-dominated component. Compared to the lower Kenilworth inner shelf deposits (i.e., Stops 1.2, 2.2, 2.3, 3.3, and 3.5), the upper Kenilworth shoreface deposits have (1) fewer Bouma-like event beds, (2) a greater amount of HCS and waverippled sandstones, (3) a higher degree of bioturbation, (4) a greater diversity of trace fossils, and (5) less carbonaceous matter. Turbidite-like beds are not present in significant amounts in the KPS3 and KPS4 inner shelf deposits. Cumulative mi (km) 54.4

(87.5)

58.8

(94.6)

Description Retrace your route. Carefully hike down the canyon and return to the vehicles. Turn right (W) onto the Price Canyon access road and drive toward Hwy 6/191. Turn right (N) onto Hwy 6/191. Return to Salt Lake City. Continue driving W-NW on

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berd een S

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Battleship Butte Blue Castle to Middle Mountain Gunnison Butte (GB) NW Mouth of Tusher Canyon Floy Wash (Subsurface Data) Salt Wash, Bootlegger Wash, Sagers Wash Pinto Wash, Nash Wash Strychnine Wash

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1 2 3 4 5 6 7 8

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L2/L1 Channel Belt Trend

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rend ine T

? e n l ine de re er ho Ab d S p n ? To sta w Lo

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Kenilworth Paleoshoreline

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Hatch Mesa Turbidites N=666

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Figure 12. (A) The paleogeography of the Western Interior Seaway during the early Campanian showing the paleoshoreline trend through the Utah Bight (modified from McGookey et al., 1972), outline of the Book Cliffs, and the location of the detailed map area in (B–D). (B) Paleogeography of the upper Aberdeen Member. The paleoshoreline trend has been “averaged” for Aberdeen parasequences (APS) APS3 to APS5 and has been projected southward along a N14°E trend (Table 2). Turbiditic-rich channel-fill deposits are present along at least two horizons (Levels 1 and 2) and are clustered in eight different areas labeled as 1–8 and shown in gray. Paleocurrent data from Gunnison Butte is shown (Table 1). (C) Inferred paleogeography of top Aberdeen Member sequence boundary. Data were collected from three outcrop localities, numbered 1–3. The Toad Stool and Hatch Mesa–Floy Wash outcrops line up along a N132°E trend (coarse-grained [CG] Layer Belt), which is similar to the mean paleocurrent trend measured from these deposits (N145°E; n = 53) (Table 1). (D) Paleogeography of the lower Kenilworth Member. The approximate downdip limit of the KPS2 upper shoreface deposits is shown (modified from Taylor and Lovell, 1995). Inner shelf turbidites are present in four areas, labeled 1–4 and shown in gray. Turbidites were transported N72°E, orthogonal to the N18°W paleoshoreline trend (Table 1).

Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior

150

(240)

200

(320)

Hwy 6 via Wellington, Price, and Soldier Summit. Take the I-15 (N) onramp at Spanish Fork and travel north on I-15 into Salt Lake City. Arrive at the Salt Palace Convention Center in Salt Lake City.

INTERPRETATION Upper Aberdeen Member A wide assortment of channel-fill deposits was examined during this field trip, including those at Bootlegger Wash (Stops 1.3 and 1.4), Gray Canyon–Tusher Canyon (Stop 2.6), and NW of Willow Bend (Stop 3.2). Many of these channel fills have previously been studied, including those at Tusher Canyon (Chan et al., 1991) and Bootlegger Wash to Sagers Wash (O’Byrne and Flint, 1993, 1995; Hampson et al., 1999). Most of these Mancos Shale–encased channel-fill deposits have similar sedimentology and sedimentary architecture. Paleocurrent data show a mean transport direction of N108°E (n = 343), which is orthogonal to the upper Aberdeen paleoshoreline trend (N14°E) (Table 1; Fig. 12B). The upper Aberdeen channel-fill deposits are interpreted as subaqueous channels that fed sediments from the delta front into the prodelta region (Fig. 13) (Pattison et al., 2005). The balance of the sedimentological evidence indicates that the channels were cut and filled in a proximal inner shelf setting by highand low-density underflows. Some of these flows appear to be linked to storm activity (i.e., low angle HCS sandstones), while others are linked to river flooding events (i.e., hyperpycnites). There is increasing evidence from studies of modern continental shelves that density underflows are common in certain inner shelf settings (Wright et al., 1988; Mulder and Syvitski, 1995; Wheatcroft et al., 1997; Mulder et al., 1998; Traykovski et al., 2000; Wheatcroft, 2000; Johnson et al., 2001; Parsons et al., 2001; Amos et al., 2003; Mulder et al., 2003; Bøe et al., 2004; Friedrichs and Wright, 2004; Hsu et al., 2004). Some of these underflows are high-density, episodic hyperpycnal flows that are storm-induced, while others are hyperpycnal flows that are longer-lived (weeks) and sustained. The former are referred to as “oceanic flood sedimentation,” while the latter are called river flood events (Wheatcroft, 2000). Both have the potential to cut scours or channels on the inner shelf, and to transport silt and fine sand into deeper water (Fig. 13; Bornhold et al., 1986; Prior et al., 1986a, 1986b; Wright et al., 1986, 1988; Ogston et al., 2000; Wheatcroft and Borgeld, 2000; Amos et al., 2003; Bøe et al., 2004). The turbiditic-rich channel-fill deposits in the upper Aberdeen Member are excellent ancient examples of oceanic and river flood deposits from an inner shelf setting. Aberdeen-Kenilworth Contact The Aberdeen-Kenilworth contact was examined at Bootlegger Wash (Stop 1.2) and Hatch Mesa (Stop 2.4), while a

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third locality was viewed with binoculars (Toad Stool, Stop 3.2). Coarse-grained deposits are rare in the Blackhawk Formation and underlying Star Point Formation (Fig. 2). Where found, these coarse-grained deposits are associated with sequence boundaries and/or transgressive surfaces of erosion. All of these examples are linked to the erosion and winnowing of shoreface attached sandstone bodies, during conditions of forced regression or transgression. Two of the best examples are transgressive lag deposits that cap the Kenilworth Member (Taylor and Lovell, 1991, 1995; Ainsworth and Pattison, 1994; Pattison, 1995) and the Panther Tongue Member (Newman and Chan, 1991; Hwang and Heller, 2002), respectively. None of these coarse-grained deposits exist as marine mudstone–encased, isolated bodies. In contrast, the coarse-grained deposits examined on this field trip are isolated pods or bodies that are completely encased in marine mudstones and are clearly detached from the highstand shoreface sandstone bodies of the Aberdeen Member (Fig. 12C). Most of these bodies have sharp bases and overlie erosional scours. Shell fragments (Baculites, Inoceramus, brachiopods, bivalves), fish teeth, and a diverse trace fossil assemblage imply fully oxygenated marine conditions. Cross bedding indicates deposition from strong traction currents. Angular to subangular mudstone clasts and moderately sorted sands suggest rapid deposition. Wood fragments, coal fragments, and finely comminuted plant matter suggest terrestrial input. The coarsening-upward grain size trend is evidence for shoaling water conditions. The uppermost lag deposit and blanketing by marine mudstones argues for transgressive ravinement of these coarse-grained bodies. This implies that the shoreface (i.e., wave base) passed through this area, both during the falling stage of sea level and the subsequent transgression. Lower Kenilworth Member The lower Kenilworth Member was examined at Bootlegger Wash (Stop 1.2), Hatch Mesa (Stops 2.2 and 2.3), Gunnison Butte (Stop 3.3), and Price River Canyon (Stop 3.5). The balance of the

Figure 13. A three-component shoreface-to-shelf depositional model (i.e., delta front, subaqueous channels, prodelta turbidites) is used to explain the inner shelf turbidite channels and lobes in the upper Aberdeen Member to lower Kenilworth Member stratigraphic interval.

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sedimentological evidence suggests that the lower Kenilworth inner shelf sandstone bodies were deposited between fair weather and storm wave base. Many of these bodies are detached from their time equivalent shoreface deposits (Fig. 12D) and consist of a wide variety of event beds, including HCS sandstones, wave-modified turbidites, classical turbidites, and hyperpycnal flow-derived turbidites (Pattison, 2004, 2005a, 2005b, 2005c; Pattison and Hoffman, 2005). Turbidite or turbidite-like event beds were likely generated by storms (i.e., oceanic floods; Wheatcroft, 2000), river flooding, high rates of sedimentation, and/or earthquakes (Walker, 1985). A three-component model is used to explain the depositional environment and setting for the lower Kenilworth Member inner shelf deposits (Fig. 13). These deposits would have been fed by a network of subaqueous channels or gullies that originated in the delta front region (Fig. 13). Similar morphological features and processes have been reported from a variety of modern continental shelf settings (Wright et al., 1988; Mulder et al., 1998; Ogston et al., 2000; Traykovski et al., 2000; Wheatcroft, 2000; Wheatcroft and Borgeld, 2000; Johnson et al., 2001; Amos et al., 2003; Mulder et al., 2003; Bøe et al., 2004; Friedrichs and Wright, 2004; Hsu et al., 2004). Sequence Stratigraphy Shelf turbidite bodies (channels and lobes) are concentrated in the upper Aberdeen Member to lower Kenilworth Member stratigraphic interval, which is subdivided into at least five parasequences. The upper Aberdeen Member is composed of three parasequences in the western part of the Book Cliffs, APS3 to APS5, and these are part of a highstand systems tract (HST) (Howell and Flint, 2003). Within the study area, the upper Aberdeen Member consists of marine mudstone–encased channel-fill successions, some of which are capped by transgressive lag deposits rich in Inoceramus and Baculites shell fragments. The lower Kenilworth Member is composed of two parasequences, KPS1 and KPS2, which are part of the overlying transgressive systems tract and early HST. Paleoshoreline trends were oriented N12°E to N16°E in the upper Aberdeen Member versus N18°W in the lower Kenilworth Member (Table 2; Fig. 12). Within the study area, turbiditic channels are more common in the upper Aberdeen Member, while turbiditic lobes are more common in the lower Kenilworth Member. The coarse-grained sandstones and lag deposits are situated at the boundary between APS5 and KPS1. The results of this work show that isolated or stray sandstone bodies are not necessarily restricted to the falling stage systems tract (FSST) or lowstand systems tract (LST). Sequence stratigraphic models should be revised to include isolated inner shelf sandstone bodies in a variety of systems tracts. CONCLUSIONS 1. Mancos Shale–encased, isolated sandstone bodies are concentrated in a relatively narrow stratigraphic interval (70– 90 m thick) within the upper Aberdeen and lower Kenilworth

members in the Green River to Thompson area, Book Cliffs, eastern Utah. 2. These isolated bodies are turbiditic-rich channels and lobes that consist of a complex mixture of wave or storm-modified turbidites, HCS sandstones, hyperpycnites, and/or classical turbidites. 3. Hyperpycnal flows were a primary mechanism for transporting fine-grained sediment from the shoreface to the inner shelf and these were likely generated during oceanic (storminduced) or river flood events. These flows also cut a network of subaqueous channels on the inner shelf. 4. Shallow marine facies models should be revised to include turbiditic-rich channels and lobes in some inner shelf settings. A three-component shoreface-to-shelf model, consisting of delta front deposits, subaqueous channels, and prodelta turbidites, is proposed to explain the depositional setting and environment of the Mancos Shale–encased sandstone bodies. 5. Sequence stratigraphic models should be revised to show that isolated sandstone bodies are not solely restricted to the FSST and LST, but can also occur in the HST. 6. Other Mancos Shale–encased isolated sandstone bodies in eastern Utah and western Colorado should be reexamined in the light of the new data and models presented herein. ACKNOWLEDGMENTS This study was funded through the Natural Sciences and Engineering Research Council of Canada Discovery Grant 238532. Bruce Ainsworth, Trevor Hoffman, Jill Stewart, Huw Williams, and Harvey Young are thanked for sharing their observations in the field. I am also grateful to Trevor for co-leading this field trip. Comments by Carol Dehler and Joel Pederson were greatly appreciated. Sharon Romanowski, Nichole Scott, and Geoff Speers are thanked for their assistance with the figures. REFERENCES CITED Ainsworth, R.B., and Pattison, S.A.J., 1994, Where have all the lowstands gone? Evidence for attached lowstand systems tracts in the Western Interior of North America: Geology, v. 22, p. 415–418, doi: 10.1130/00917613(1994)022<0415:WHATLG>2.3.CO;2. Amos, C.L., Li, M.Z., Chiocci, F.L., La Monica, G.B., Cappucci, S., King, E.H., and Corbani, F., 2003, Origin of shore-normal channels from the shoreface of Sable Island, Canada: Journal of Geophysical Research, v. 108 (C3), p. 39-1–39-16. Arnott, R.W., and Southard, J.B., 1990, Exploratory flow-duct experiments on combined-flow bed configurations, and some implications for interpreting storm-event stratification: Journal of Sedimentary Petrology, v. 60, p. 211–219. Balsley, J.K., 1980, Cretaceous wave-dominated delta systems, Book Cliffs, east-central Utah: American Association of Petroleum Geologists, Continuing Education Course, Field Guide, 163 p. Bøe, R., Bugge, T., Rise, L., Eidnes, G., Eide, A., and Mauring, E., 2004, Erosional channel incision and the origin of large sediment waves in Trondheimsfjorden, central Norway: Geo-Marine Letters, v. 24, p. 225–240, doi: 10.1007/s00367-004-0180-3. Bornhold, B.D., Yang, Z.-S., Keller, G.H., Prior, D.B., Wiseman, W.J., Jr., Wang, Q., Wright, L.D., Xu, W.D., and Zhuang, Z.Y., 1986, Sedimentary framework of the modern Huanghe (Yellow River) delta: Geo-Marine Letters, v. 6, p. 77–83.

Recognition and interpretation of isolated shelf turbidite bodies in the Cretaceous Western Interior Chan, M.A., 1992, Oolitic ironstone of the Cretaceous Western Interior Seaway, east-central Utah: Journal of Sedimentary Petrology, v. 62, p. 693–705. Chan, M.A., Newman, S.L., and May, F.E., 1991, Deltaic and shelf deposits in the Cretaceous Blackhawk Formation and Mancos Shale, Grand County, Utah: Utah Geological Survey Miscellaneous Publication 91-6, 83 p. Cole, R.D., and Young, R.G., 1991, Facies characterization and architecture of a muddy shelf-sandstone complex: Mancos B interval of Upper Cretaceous Mancos Shale, northwest Colorado–northeast Utah, in Miall, A.D., and Tyler, N., eds., The Three-dimensional facies architecture of terrigenous clastic sediments and its implications for hydrocarbon discovery and recovery: Society for Sedimentary Geology (SEPM) Concepts in Sedimentology and Paleontology, v. 3, p. 277–287. Cole, R.D., Young, R.G., and Willis, G.C., 1997, The Prairie Canyon Member, a new unit of the Upper Cretaceous Mancos Shale, west-central Colorado and east-central Utah: Utah Geological Survey Miscellaneous Publication 97-4, 23 p. Frey, R.W., and Howard, J.D., 1990, Trace fossils and depositional sequences in a clastic shelf setting, Upper Cretaceous of Utah: Journal of Paleontology, v. 64, p. 803–820. Friedrichs, C.T., and Wright, L.D., 2004, Gravity-driven sediment transport on the continental shelf: implications for equilibrium profiles near river mouths: Coastal Engineering, v. 51, p. 795–811, doi: 10.1016/j.coastale ng.2004.07.010. Hampson, G.J., 2000, Discontinuity surfaces, clinoforms, and facies architecture in a wave-dominated, shoreface-shelf parasequence: Journal of Sedimentary Research, v. 70, p. 325–340. Hampson, G.J., 2004, Day 3 Mancos B sandstones: Grand Junction, Colorado, Society for Sedimentary Geology (SEPM) Research Conference Field Guide, Recent Advances in Shoreline-Shelf Stratigraphy, 61 p. Hampson, G.J., Burgess, P.M., and Howell, J.A., 2001, Shoreface tongue geometry constrains history of relative sea-level fall: Examples from Late Cretaceous strata in the Book Cliffs, Utah: Terra Nova, v. 13, p. 188–196, doi: 10.1046/j.1365-3121.2001.00337.x. Hampson, G.J., Howell, J.A., and Flint, S.S., 1999, A sedimentological and sequence stratigraphic re-interpretation of the Upper Cretaceous Prairie Canyon Member (“Mancos B”) and associated strata, Book Cliffs area, Utah, U.S.A.: Journal of Sedimentary Research, v. 69, p. 414–433. Hettinger, R.D., and Kirschbaum, M.A., 2002, Stratigraphy of the Upper Cretaceous Mancos Shale (upper part) and Mesaverde Group in the southern part of the Uinta and Piceance Basins, Utah and Colorado: U.S. Geological Survey Geological Investigations I-2764, 21 p. Howell, J.A., and Flint, S.S., 2003, 9 Sequences and systems tracts in the Book Cliffs, and 10 Sequence stratigraphical evolution of the Book Cliffs succession, in Coe, A.L., ed., The sedimentary record of sea-level change: Cambridge University Press, p. 179–208. Hsu, S.-C., Lin, F.-J., Jeng, W.-L., Chung, Y., Shaw, L.-M., and Hung, K.W., 2004, Observed sediment fluxes in the southwesternmost Okinawa Trough enhanced by episodic events: Flood runoff from Taiwan rivers and large earthquakes: Deep-Sea Research. Part I: Oceanographic Research Papers, v. 51, p. 979–997, doi: 10.1016/j.dsr.2004.01.009. Hwang, I.-G., and Heller, P.L., 2002, Anatomy of a transgressive lag: Panther Tongue Sandstone, Star Point Formation, central Utah: Sedimentology, v. 49, p. 977–999, doi: 10.1046/j.1365-3091.2002.00486.x. Johnson, K.S., Paull, C.K., Barry, J.P., and Chavez, F.P., 2001, A decadal record of underflows from a coastal river into the deep sea: Geology, v. 29, p. 1019– 1022, doi: 10.1130/0091-7613(2001)029<1019:ADROUF>2.0.CO;2. Kellogg, H.E., 1977, Geology and petroleum of the Mancos B Formation, Douglas Creek Arch area Colorado and Utah, in Veal, H.K., ed., Exploration frontiers of the central and southern Rockies: Denver, Colorado, Rocky Mountain Association of Geologists, 1977 Symposium, p. 167–179. McGookey, D.P., Haun, J.D., Hale, L.A., Goodell, H.G., McCubbin, D.G., Weimer, R.J., and Wulf, G.R., 1972, Cretaceous systems, in Mallory, W.W., ed., Geologic atlas of the Rocky Mountain region: Rocky Mountain Association of Geologists, p. 190–228. Mulder, T., and Syvitski, J.P.M., 1995, Turbidity currents generated at river mouths during exceptional discharges to the world oceans: Journal of Geology, v. 103, p. 285–299. Mulder, T., Syvitski, J.P.M., and Skene, K.I., 1998, Modeling of erosion and deposition by turbidity currents generated at river mouths: Journal of Sedimentary Research, v. 68, p. 124–137. Mulder, T., Syvitski, J.P.M., Migeon, S., Faugères, J.-C., and Savoye, B., 2003, Marine hyperpycnal flows: initiation, behaviour and related deposits: A

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Cliffs, Utah, USA: Sedimentary Geology, v. 177, p. 131–144, doi: 10.1016/j.sedgeo.2005.02.005. Pattison, S.A.J., 2005c, A stray sandstone finds a home: time equivalency of the Hatch Mesa succession and the lower Kenilworth Member, Book Cliffs, Utah: Geological Society of America Abstracts with Programs, v. 37, no. 6, p. 44. Pattison, S.A.J., 2005d, Tectonic significance of shelf turbidite bodies in the upper Aberdeen Member to lower Kenilworth Member stratigraphic interval, Cordilleran Foreland Basin, Book Cliffs, eastern Utah: Calgary, Alberta, American Association of Petroleum Geologists Annual Convention, Abstract, p. A106. Pattison, S.A.J., and Hoffman, T.A., 2005, Evidence of fine-grained sediment transport by turbidity currents, hyperpycnal flows and storm waves: inner shelf prodelta turbidite complex, Hatch Mesa succession, Book Cliffs, Utah: Calgary, Alberta, American Association of Petroleum Geologists Annual Convention, Abstract, p. A106. Pattison, S.A.J., Hoffman, T.A., and Ainsworth, R.B., 2005, Prodelta turbiditic channels and lobes, upper Aberdeen Member to lower Kenilworth Member, Green River Embayment, Book Cliffs, Utah: Integral components of river-dominated deltas: Calgary, Alberta, American Association of Petroleum Geologists Annual Convention, Abstract, p. A106. Pemberton, S.G., Spila, M., Pulham, A.J., Saunders, T., MacEachern, J.A., Robbins, D., and Sinclair, I.K., 2001, Ichnology and sedimentology of shallow to marginal marine systems: Ben Nevis and Avalon reservoirs, Jeanne d’Arc Basin: Geological Association of Canada, Short Course Notes, v. 15, 343 p. Posamentier, H.W., 2001, Lowstand alluvial bypass systems: incised vs. unincised: AAPG Bulletin, v. 85, p. 1771–1793. Prior, D.B., Yang, Z.-S., Bornhold, B.D., Keller, G.H., Lin, Z.H., Wiseman, W.J., Jr., Wright, L.D., and Lin, T.C., 1986a, The subaqueous delta of the modern Huanghe (Yellow River): Geo-Marine Letters, v. 6, p. 67–75. Prior, D.B., Yang, Z.-S., Bornhold, B.D., Keller, G.H., Lu, N.Z., Wiseman, W.J., Jr., Wright, L.D., and Zhang, J., 1986b, Active slope failure, sediment collapse, and silt flows on the modern subaqueous Huanghe (Yellow River) delta: Geo-Marine Letters, v. 6, p. 85–95. Saunders, T., and Pemberton, S.G., 1986, Trace fossils and sedimentology of the Appaloosa Sandstone: Bearpaw–Horseshoe Canyon Formation transition, Dorothy, Alberta: Canadian Society of Petroleum Geologists Field Trip Guide Book, 117 p. Spieker, E.M., 1931, The Wasatch Plateau coal fields, Utah: U.S. Geological Survey Bulletin 819, 210 p. Stevens, K.M., and Chaiwongsaen, N., 2003, Evaluating the origin of isolated Cretaceous sandstones encased in marine mudstone: The Prairie Canyon Member of the Mancos Shale, Hatch Mesa, east-central Utah: Salt Lake City, Utah, American Association of Petroleum Geologists Annual Convention, Abstract, p. A163. Stevens, K.M., Anderson, D.S., Gardner, M.H., and Wagner, J.B., 2005, Reevaluating the origin of isolated sandstones encased in marine mudstone: the Cretaceous Prairie Canyon Member of the Mancos Shale at Hatch Mesa, east-central Utah: Calgary, Alberta, American Association of Petroleum Geologists Annual Convention, Abstract, p. A133. Swift, D.J.P., Hudelson, P.M., Brenner, R.L., and Thompson, P., 1987, Shelf construction in a foreland basin: Storm beds, shelf sandbodies, and shelfslope depositional sequences in the Upper Cretaceous Mesaverde Group, Book Cliffs, Utah: Sedimentology, v. 34, p. 423–457. Taylor, D.R., and Lovell, R.W.W., 1991, Recognition of high-frequency sequences in the Kenilworth Member of the Blackhawk Formation, Book Cliffs, Utah, in Van Wagoner, J.C., Nummedal, D., Jones, C.R., Taylor, D.R., Jennette, D.C., and Riley, G.W., eds., Sequence stratigraphy applications to shelf

sandstone reservoirs: Outcrop to subsurface examples: American Association of Petroleum Geologists, Field Conference Guidebook. Taylor, D.R., and Lovell, R.W.W., 1995, High-frequency sequence stratigraphy and paleogeography of the Kenilworth Member, Blackhawk Formation, Book Cliffs, Utah, U.S.A., in Van Wagoner, J.C., and Bertram, G.T., eds., Sequence stratigraphy of foreland basin deposits—outcrop and subsurface examples from the Cretaceous of North America: American Association of Petroleum Geologists Memoir 64, p. 257–275. Taylor, K.G., Gawthorpe, R.B., Curtis, C.D., Marshall, J.M., and Awwiller, D.N., 2000, Carbonate cementation in a sequence-stratigraphic framework: Upper Cretaceous sandstones, Book Cliffs, Utah-Colorado: Journal of Sedimentary Research, v. 70, p. 360–372. Traykovski, P., Geyer, W.R., Irish, J.D., and Lynch, J.F., 2000, The role of wave-induced density-driven fluid mud flows for cross-shelf transport on the Eel River continental shelf: Continental Shelf Research, v. 20, p. 2113–2140, doi: 10.1016/S0278-4343(00)00071-6. Van Wagoner, J.C., 1991, Sequence stratigraphy and facies architecture of the Desert Member of the Blackhawk Formation and the Castlegate Formation in the Book Cliffs of eastern Utah and western Colorado, in Van Wagoner, J.C., Nummedal, D., Jones, C.R., Taylor, D.R., Jennette, D.C., and Riley, G.W., eds., Sequence stratigraphy applications to shelf sandstone reservoirs—outcrop to subsurface examples: American Association of Petroleum Geologists, Field Conference Guidebook. Van Wagoner, J.C., 1995, Sequence stratigraphy and marine to nonmarine facies architecture of foreland basin strata, Book Cliffs, Utah, U.S.A., in Van Wagoner, J.C., and Bertram, G.T., eds., Sequence stratigraphy of foreland basin deposits—outcrop and subsurface examples from the Cretaceous of North America: American Association of Petroleum Geologists Memoir 64, p. 137–223. Van Wagoner, J.C., Mitchum, R.M., Campion, K.M., and Rahmanian, V.D., 1990, Siliciclastic sequence stratigraphy in well logs, cores, and outcrops: American Association of Petroleum Geologists, Methods in Exploration Series 7, 55 p. Walker, R.G., 1985, Geological evidence for storm transportation and deposition on ancient shelves, in Tillman, R.W., Swift, D.J.P., and Walker, R.G., eds., Shelf sands and sandstone reservoirs: Society of Economic Paleontologists and Mineralogists, Short Course Notes 13, p. 243–302. Wheatcroft, R.A., 2000, Oceanic flood sedimentation: A new perspective: Continental Shelf Research, v. 20, p. 2059–2066, doi: 10.1016/S02784343(00)00062-5. Wheatcroft, R.A., and Borgeld, J.C., 2000, Oceanic flood deposits on the northern California shelf: large-scale distribution and small-scale physical properties: Continental Shelf Research, v. 20, p. 2163–2190, doi: 10.1016/S0278-4343(00)00066-2. Wheatcroft, R.A., Sommerfield, C.K., Drake, D.E., Borgeld, J.C., and Nittrouer, C.A., 1997, Rapid and widespread dispersal of flood sediment on the northern California margin: Geology, v. 25, p. 163–166, doi: 10.1130/ 0091-7613(1997)025<0163:RAWDOF>2.3.CO;2. Wright, L.D., Yang, Z.-S., Bornhold, B.D., Keller, G.H., Prior, D.B., and Wiseman, W.J., Jr., 1986, Hyperpycnal plumes and plume fronts over the Huanghe (Yellow River) delta front: Geo-Marine Letters, v. 6, p. 97–105. Wright, L.D., Wiseman, W.J., Bornhold, B.D., Prior, D.B., Suhayda, J.N., Keller, G.H., Yang, Z.-S., and Fan, Y.B., 1988, Marine dispersal and deposition of Yellow River silts by gravity-driven underflows: Nature, v. 332, p. 629–632, doi: 10.1038/332629a0. Young, R.G., 1955, Sedimentary facies and intertonguing in the Upper Cretaceous of the Book Cliffs, Utah-Colorado: Geological Society of America Bulletin, v. 66, p. 177–202.

Printed in the USA

Geological Society of America Field Guide 6 2005

Geologic hazards of the Wasatch Front, Utah Barry J. Solomon Francis X. Ashland Richard E. Giraud Michael D. Hylland Utah Geological Survey, P.O. Box 146100, Salt Lake City, Utah 84114-6100, USA Bill D. Black Western GeoLogic, LLC, 74 N Street, Salt Lake City, Utah 84103-3860, USA Richard L. Ford Michael W. Hernandez Department of Geosciences, Weber State University, 2507 University Circle, Ogden, Utah 84408-2505, USA David H. Hart Capitol Preservation Board, 110 Senate Building, Salt Lake City, Utah 84114, USA

ABSTRACT The Wasatch Front, Utah’s population center, faces threats from several geologic hazards. These range from hazards that occur somewhere in the region almost every year, such as landslides and debris flows, to potentially catastrophic but infrequent hazards resulting from large earthquakes along the Wasatch fault zone. Study of these hazards is an ongoing process, and recent related research will be discussed on this field trip. We will observe an active landslide in Salt Lake City and hazard-reduction techniques implemented following fire-related debris flows near Farmington. We will examine evidence of recent flooding from increased levels of Great Salt Lake and discuss flooding hazards posed by the lake. We will observe the effects of large prehistoric earthquakes along the Wasatch fault zone, the longest active, normal-slip fault zone in the United States, and will discuss paleoseismology of the fault zone and the potential for earthquake ground shaking in the Salt Lake Valley. We will also examine ongoing efforts to seismically retrofit the Utah State Capitol to withstand strong earthquake ground shaking while preserving the historical integrity of its architecture. Keywords: geologic hazards, landslide, earthquake, debris flow, flooding, Wasatch fault zone, Wasatch Front. INTRODUCTION The Wasatch Range extends more than 100 mi (160 km) through north-central Utah. The bulk of Utah’s population resides in rapidly growing metropolitan areas along the range

front, and the Wasatch Front is subject to a variety of geologic hazards due to a unique combination of geologic, topographic, and climatic conditions. The Wasatch Front occupies a series of north-trending valleys at the foot of the western slope of the Wasatch Range. The

Solomon, B.J., Ashland, F.X., Giraud, R.E., Hylland, M.D., Black, B.D., Ford, R.L., Hernandez, M.W., and Hart, D.H., 2005, Geologic hazards of the Wasatch Front, Utah, in Pederson, J., and Dehler, C.M., eds., Interior Western United States: Geological Society of America Field Guide 6, p. 505–524, doi: 10.1130/ 2005.fld006(22). For permission to copy, contact [email protected]. © 2005 Geological Society of America

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Figure 1. Field trip route (heavy line) along the central Wasatch Front, Utah, and numbered stops.

mountains rise steeply as much as 7100 ft (2200 m) above valley floors, reaching elevations near 12,000 ft (3700 m) above sea level. This impressive relief is the result of ongoing displacement along the Wasatch fault zone. The fault zone separates the Basin and Range Province to the west from the middle Rocky Mountains to the east and is a major intraplate tectonic boundary. The Wasatch fault zone is the longest active normal-slip fault zone in the United States and one of several fault zones in the region considered capable of producing large (M > 7) earthquakes. These earthquakes would generate surface fault rupture and strong ground shaking, perhaps accompanied by seismically induced liquefaction and landslides. The central five of the 10 independent segments of the Wasatch fault zone are considered the most active, having produced at least 16 large earthquakes in the past 6000 yr (McCalpin and Nishenko, 1996). Large earthquakes on the central five segments recur, on average, about every 350 yr, with the last occurring ca. 620 yr ago. Great Salt Lake, a remnant of the much larger Pleistocene Lake Bonneville, forms the western boundary of the northern Wasatch Front. The well-established chronology of the four major shorelines of the Bonneville Lake cycle is often used to determine the age of movement for landslides and surface faulting along the Wasatch Front, and Lake Bonneville deposits

and shorelines have a profound influence on the distribution of Wasatch Front geologic hazards. For example, loose, saturated Bonneville sands are potentially liquefiable, and steep cliffs near the erosion platforms of Bonneville shorelines increase the potential for slope failures in lake sediment. Thick deposits of soft, fine-grained lacustrine sediment from Lake Bonneville and older Pleistocene deepwater lakes that inundated the valleys could amplify earthquake ground motions. Post-Bonneville Great Salt Lake, occupying a closed basin within the internally draining Great Basin, is subject to climate-induced fluctuations. Rising lake levels between 1983 and 1987 caused flooding in areas along and near the gently sloping lake shores. In the winter, frontal storms traveling east from the Pacific Ocean encounter the Wasatch Range and produce heavy snowfall in the mountains. Snow avalanches are common and present a significant, widespread hazard. Freeze-thaw cycles in steep exposures of fractured rock produce rock falls. Rapid melting of a lingering snowpack periodically results in slope failures, debris flows, and stream and alluvial fan flooding. Convective storms in the spring and late summer also contribute to these hazards. This field trip (Fig. 1) provides an opportunity to observe and discuss several of the most significant types of geologic hazards of the Wasatch Front and examine the results of recent and ongoing related research. Field trip topics include (1) slow, active landsliding near the mouth of City Creek Canyon in Salt Lake City; (2) site conditions, earthquake ground shaking, and the Utah State Capitol seismic retrofit; (3) prehistoric liquefactioninduced landsliding in lacustrine sediments near Farmington; (4) debris-flood and debris-flow hazard reduction techniques on alluvial fans at the base of the Wasatch Range in Davis County; (5) Pleistocene Lake Bonneville shorelines and evidence of historical Great Salt Lake flooding on Antelope Island; (6) faulted Pleistocene (?) deposits in western Salt Lake Valley; and (7) surface fault rupture and paleoseismology on the active Salt Lake City segment of the Wasatch fault zone. FIELD TRIP Directions to Stop 1 From the Salt Palace Convention Center at 100 South West Temple in Salt Lake City (Fig. 1), proceed south on West Temple. Turn left on 200 South and after two blocks turn left on State Street. Proceed north on State Street ~0.3 mi (0.5 km), and turn right on South Temple, continuing ~0.3 mi (0.5 km). Turn left on B Street and proceed ~0.8 mi (1.3 km) to 11th Ave. Continue straight ahead, driving north on East Bonneville Boulevard ~0.8 mi (1.3 km) and park before crossing City Creek. Stop 1—East Capitol Boulevard–City Creek Landslide, Salt Lake City The East Capitol Boulevard–City Creek (CBCC) landslide is possibly the best example of a recurrently active landslide

Geologic hazards of the Wasatch Front, Utah along the Wasatch Front. The landslide predates the earliest (1937) aerial photographs of the Salt Lake City area (Van Horn et al., 1972). Movement occurred in five of the seven years between 1998 and 2004 (inclusive). About 10.8 ft (3.3 m) of movement occurred in a single year (2002), surprisingly the driest calendar year of a drought that lasted between 1999 and 2004 (Fig. 2). Retrogressive enlargement of the landslide in 1998 damaged parts of a backyard and threatened a house above the western part of the slide. A drilled pier (caisson) wall was installed by early 1999 to protect the remainder of the property. Offset on the main scarp subsequently continued, requiring stabilization of the scarp face with soil nails and shotcrete in 2004. Minor offset on the main scarp postdates the shotcrete application. At the northern corner of the landslide, continued offset of the main scarp oversteepens the slope and threatens a tennis court. Progressive enlargement of the landslide, which began in 1999 with the formation of a new frontal toe thrust, and downslope movement of slide debris have resulted in encroachment on an inlet structure that drains a perennial creek that flows along the left (SE) flank of the slide into City Creek. Research by the Utah Geological Survey (UGS) on landslide movement and precipitation (Ashland, 2003) defined a method for predicting movement of recurrently active landslides based on recognition of cumulative instability-threshold precipitation levels. This method tracks both antecedent precipitation and conditions at the onset of the snowmelt that overlap with triggering of most northern Utah landslides. Geology The CBCC landslide is on a southeast-facing slope above a tributary drainage that flows into City Creek. The drainage was once ephemeral (Dames & Moore, 1979), but now flows year round. Late Pleistocene Lake Bonneville fine-grained sediments underlie the head and crown of the landslide (Dames & Moore, 1979, 1981; Personius and Scott, 1992). Exposures in the main scarp consist mostly of weakly laminated silt. Test pits and boreholes in the upper part and the crown area of the landslide indicate that soils consist primarily of interbedded silt, silty sand, sand, and minor gravel (Dames & Moore, 1979, 1981). The lacustrine sediments in turn overlie soils derived from Tertiary (Paleogene) sedimentary and volcanic rocks (Personius and Scott, 1992). Locally, coarse-grained fill overlies natural soils near the crest of the slope and head of the slide. Landslide Description The landslide is funnel shaped in plan view, narrowing in width downslope. At its head, the landslide is nearly 400 ft (122 m) wide, but it is only ~45 ft (14 m) wide at the toe. In 1979, the landslide was ~480 ft (146 m) long, ~360 ft (110 m) wide at the head, and ~160 ft (49 m) wide at its lower part upslope of the toe. By 1998, the landslide area was ~19,000 yd2 (16,000 m2) with an estimated volume between 130,000 and 240,000 yd3 (99,000–148,000 m3). Profiling of the landslide in 1999 indi-

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Figure 2. Displacement history of the toe of the East Capitol Blvd– City Creek landslide, June 1998 to December 2004. Inset shows toe displacement in 2004 in detail. Note differences in scale of y-axes.

cated that it was ~570–580 ft (174–177 m) long, or ~90–100 ft (28–31 m) longer than in 1979. The difference in length suggests an average annual rate of stretching of ~4.5–5 ft/yr (1.4–1.5 m/ yr) during the intervening 20 yr. Survey data document ~29.5 ft (9 m) of stretching in the upper part of the landslide between May 1987 and April 2001. Movement History Between 1998 and 2004, the landslide moved in all but two years (2000 and 2003). Total movement of the landslide toe between June 1998 and December 2004 exceeded 23.6 ft (7.2 m). Movement measurements by the UGS at the toe of the landslide began on 5 June 1998. Although complete movement data are unavailable for 1998, the landslide moved a little over 8 ft (2.4 m) between 5 June 1998, and 5 June 1999 (Fig. 2). Movement typically triggers in early March and suspends between late May and early July (Ashland, 2002) (see inset in Fig. 2). Wetter than normal conditions in late spring or summer can lengthen the duration of movement or cause reactivation of the slide in the latter part of the year. Instability-Threshold Precipitation Level Movement of many landslides in northern Utah is triggered by a transient rise in groundwater levels or accumulation of ephemeral perched groundwater associated with the snowmelt. Groundwater level monitoring since 1999 has documented that natural seasonal peak groundwater levels are significantly reduced by the lack of a late winter snowpack (Ashland et al., 2005). Cumulative precipitation is a reasonable basis for estimating the fluctuation in landslide groundwater levels. Ashland (2003) documented that groundwater levels along the Wasatch Front rise during extended long-term periods of excess pre-

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cipitation. In addition, Ashland (2003) showed a correlation between cumulative precipitation and landslide movement suggesting that groundwater levels fluctuate with the cumulative precipitation budget (excess versus deficit precipitation conditions) of a site. Figure 3 shows mean cumulative-precipitation curves for active and inactive periods between 1998 and 2004 that define the limits on the instability-threshold precipitation levels for the CBCC landslide. The instability-threshold precipitation levels fall between the two curves and are inferred to be near the normal cumulative-precipitation curve. The shaded area indicates the critical period for landslide movement (typically March through May in northern Utah). Infiltration of cumulative excess precipitation prior to and during the critical period causes a rise in groundwater levels relative to the previous year. The exceptionally low instability-threshold precipitation levels of the CBCC landslide suggest that human modifications of the hillslope have significantly reduced the natural stability of the slide. The most important change that impacts slope stability may be a long-term rise in the base groundwater level due to infiltration of landscape irrigation water (lawn watering) in the residential subdivision directly upslope of the landslide and redirected runoff from the residential lots onto the slide. As a result of the elevated base groundwater levels, the transient snowmeltinduced groundwater-level rise needed to reach an instabilitythreshold level and trigger movement has likely decreased in magnitude in the past few decades, so that it occurs even in years with near-normal cumulative precipitation. Table 1 shows that a prerequisite for renewed landslide movement of the CBCC landslide is cumulative excess precipitation sometime during the critical period. In the only two years in which the landslide remained dormant (2000 and 2003), cumulative deficit conditions existed during the entire critical period. In 1999, cumulative-excess precipitation conditions occurred only in May. Most (~90%) of the movement at the toe of the landslide in 1999 occurred in May and early June, during the transition from cumulative-deficit to cumulative-excess precipitation conditions. Total movement at the toe of the landslide shows a roughly inverse correlation with the maximum cumulative excess pre-

Figure 3. Calibrated cumulative-precipitation curves defining instability-threshold precipitation levels for the East Capitol Blvd–City Creek landslide. Cumulative precipitation is shown for a landslide water year (LWY) beginning 1 September. The critical period during which landslide movement typically triggers is 1 March–1 June (shaded box). Instability-threshold precipitation levels fall somewhere between the mean cumulative precipitation for LWYs in which movement did (black curve) and did not (gray curve) occur and is inferred to be near the normal precipitation (not shown). Cumulative precipitation for the 1997–1998 LWY (dashed line), shown for comparison, remained greater after 1 March than mean cumulative precipitation for LWYs in which movement occurred.

cipitation during the critical period. The largest measured annual displacements at the toe occurred at the end of successive years in which excess precipitation existed in the critical period (1999 and 2001). The cumulative effects of successive wetter-than-normal years may have more influence on groundwater levels in a landslide and total annual displacement than the excess precipitation in a single year (such as in the 2003–2004 landslide water year). In addition, other factors such as landslide geometry and boundary conditions may have an increasing influence on the magnitude of total annual displacements, particularly near the end of a period of prolonged movement (1998–2004).

TABLE 1. SUMMARY OF CUMULATIVE-PRECIPITATION CONDITIONS NECESSARY TO TRIGGER RENEWED MOVEMENT Cumulative-precipitation budget (inches) Months Sept.–Feb. Sept.–Mar. Sept.–Apr. Sept.–May Active?

Landslide water year 97-98 3.65 4.71 4.68 3.92

98-99 –0.18 –1.29 –0.32 0.40

99-00 –0.77 –1.84 –3.20 –3.38

00-01 0.94 0.58 0.92 –0.66

01-02 –0.51 0.05 0.42 –0.90

02-03 –3.04 –3.56 –4.03 –4.16

03-04 2.62 1.59 1.85 1.00

Yes

Yes

No

Yes

Yes

No

Yes

Note: Bold indicates cumulative excess conditions; italics indicate near normal conditions; normal print indicates cumulative deficit conditions.

Geologic hazards of the Wasatch Front, Utah Directions to Stop 2 Cross City Creek and proceed south on West Bonneville Blvd for ~0.8 mi (1.3 km). Turn right on 500 North, continue for ~0.3 mi (0.5 km), and turn left on Columbus Street. After ~0.2 mi (0.3 km), turn left into the State Capitol parking lot. Stop 2—Site Conditions, Earthquake Ground Shaking, and the Utah State Capitol Seismic Retrofit Site Conditions and Earthquake Ground Shaking in Salt Lake Valley Unconsolidated surficial Quaternary deposits in Salt Lake Valley were mostly derived from late Pleistocene Lake Bonneville, but surficial and underlying deposits also include pre-Bonneville alluvial fan deposits, late Pleistocene (and older?) glacial deposits (till and outwash), and Holocene stream alluvium, alluvial fan deposits, and lacustrine and deltaic deposits. Sediments on the northeast margin of Salt Lake Valley occupy the footwall of the active Wasatch fault zone and overlie shallow, older (pre-Bonneville) semi-consolidated valley fill sediments, and/or weathered Tertiary or older rock. Unconsolidated deposits on the west side of Salt Lake Valley possibly overlie a relatively shallow rock bench that extends several miles from the base of the Oquirrh Mountains, where Holocene faulting is absent. Site conditions reflect the properties of these near-surface geologic materials and can have a significant impact on earthquake ground shaking. Analytical (Martin and Dobry, 1994) and empirical ground motion studies (Borcherdt, 1994) suggest site conditions can be reasonably characterized using the average shear-wave velocity in the upper 100 ft (30 m) (Vs30). The International Building Code (IBC), adopted in Utah in 2002 and updated with a 2003 edition (International Code Council, 2002), uses Vs30 to group sites into five broad site classes that are assigned factors for use in estimating earthquake loads in the design of engineered buildings and structures. The five site classes are designated A through E, and are respectively named hard rock, rock, very dense soil and soft rock, stiff soil profile, and soft soil profile; a sixth IBC site class, F, includes soil that poses engineering problems for building foundations and is not characterized by Vs30. Site conditions in Salt Lake Valley were mapped by combining 30 new shear-wave velocity profiles with the areal limited, existing shear-wave velocity profile data and previous mapping of Ashland and McDonald (2003). The existing site conditions map was simplified by redefining some previously mapped unit and subunit boundaries and characterized all four previously mapped Quaternary (unconsolidated) site condition units. Some shear-wave velocity profiles encountered buried rock, providing better characterization of Vs30 in the three previously mapped rock site condition units in the mountain areas bordering Salt Lake Valley (Ashland and McDonald, 2003) and of shallow impedance contrasts along local valley margins.

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Figure 4 shows the newly revised site conditions map for Quaternary units in Salt Lake Valley, and Table 2 summarizes Vs30 in these units. The Vs30 statistics in Table 2 allow for estimating the potential variation in Vs30 at a site. IBC site classes in Salt Lake Valley Quaternary deposits range from site class E to site class C. A boundary between IBC site classes falls within the range of Vs30 in all but one of the Quaternary site conditions units. The Utah State Capitol lies within site conditions unit Q02 (Fig. 4), which includes Vs30 values characteristic of IBC site classes D and C. IBC site coefficients for these site classes indicate the potential for significant amplification of strong earthquake ground motions, and such motions may be particularly strong because the Capitol lies near the Wasatch fault zone. Utah State Capitol Seismic Retrofit In 1888, Salt Lake City donated land to the State of Utah for the construction of the State Capitol at its present site. Construction began in December 1912, and its cornerstone was set in 1914. The legislature relocated to the new Capitol in February 1915 and in September 2004, the building was closed for seismic renovations (Cooper Roberts Simonson Architecture, 2000). When constructed, the Utah State Capitol was one of the first concrete frame buildings west of the Mississippi River. While this technological advancement was new and exciting, it did not replace many of the common construction practices of the

Figure 4. Site conditions map of Salt Lake Valley, Utah (revised from Ashland and McDonald, 2003). The Weber segment of the Wasatch fault zone (WFZ) (discussed at Stop 3) lies to the north of Salt Lake Valley. For explanation of site condition units (Q01 through Q04), see Table 2.

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TABLE 2. SUMMARY OF Vs30 FOR SALT LAKE VALLEY QUATERNARY SITE CONDITION UNITS Unit*

Mean Vs30 (m/sec)

Standard deviation (%)

Maximum (m/sec)

Minimum (m/sec)

Median (m/sec)

Q01 Q02 Q03 Q04

197† 293 407 456

19† 18 25 7

325 469 708 510

151 212 294 413

188† 288 394 459

IBC site class Range in site Number of Vs (mean) class profiles D D C C

E to D D to C D to C C

66 40 19 6

*Q01—Lacustrine and alluvial silt and clay; Q02—Lacustrine sand, silt, and clay; Q03—Lacustrine and alluvial gravel; Q04—Pre-Bonneville alluvial fan deposits. † Revised from Ashland and McDonald (2003).

day. Those techniques did not account for the strong earthquake ground shaking that recent studies indicate may occur along the Wasatch Front. The exterior of the building was made of large, cut granitic stone from Little Cottonwood Canyon, ~20 mi (30 km) south of the Capitol. Rather than anchoring the stone to the frame, as is current practice, the stone blocks were stacked one on top of another, as a load-bearing masonry wall would have been built at the time. Use of a concrete frame was a new technique and the need for reinforcing bars was not well understood. Some steel was added to the frame to provide additional strength, but sometimes the steel was used only as a form-tie element to hold the formwork in place. Despite older construction techniques, much of the Capitol was quite strong and able to withstand significant vertical force. However, studies have indicated that when the structure is subjected to a lateral load of a large M 7.0 earthquake, the building may fail and collapse at the parapet level. An important architectural feature of the Utah State Capitol, the dome-and-drum assembly on top of the building, was in very poor condition. Engineers estimated that a moderate earthquake could cause its collapse because of the reduced strength of the concrete used in its construction (Reaveley Engineers and Associates, 2003). Compression tests of concrete core samples from the drum indicated that concrete breaks could occur at stresses as low as 900 psi (6000 kPa) (AGRA Earth and Environmental, 2000), perhaps because the concrete was poured during the winter and may have frozen during the cold evening hours. The concrete also had a very high water-to-cement ratio (AGRA Earth and Environmental, 2000), further contributing to its low strength. Because of the nature of the concrete, outdated construction techniques, and potential amplification of earthquake ground shaking by the tall dome-and-drum assembly (Reaveley Engineers and Associates, 2003), the seismic retrofit would have to provide a major reduction in force displacement.

level of demand on the structural engineer and the design team to provide the needed seismic design without altering the building’s architectural integrity. After studying various structural systems and their impacts, the structural engineers recommended that a combination of base isolation and shear walls be used. Base isolation would decouple the building from the horizontal component of earthquake ground motions, and shear walls would resist and limit interstory drift. Base isolation also reduced requirements for seismic reinforcement, avoiding damage to the integrity of the architecture of the Capitol (Reaveley Engineers and Associates, 2003). Base isolation will lengthen the periodic response of the Capitol from <~1 s to 3 or 4 s (Fig. 5). The shorter response results in lateral accelerations of ~1.4 g, whereas the longer response resulting from base isolation lowers lateral accelerations to ~0.3 g. This reduction allows the use of other seismic design elements that are less intrusive than would otherwise be required to resist the effects of greater accelerations.

Seismic Base Isolation and Shear Walls The seismic design of the Capitol had to account for the seismicity of the site, the safety of building occupants, and the historic and symbolic nature of the structure. This placed a higher

Figure 5. Typical seismic-response spectrum illustrating the effects of base isolation and shear-wall reinforcement (Reaveley Engineers and Associates, 2003). Base isolation lengthens the periodic response of the Capitol, decoupling the building from the horizontal component of ground shaking. Shear walls resist and limit inter-story drift.

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To design the base isolation system, site-specific data were collected from several boreholes drilled to depths of at least 350 ft (110 m). Data included the results of downhole seismic surveys, which were used to develop several synthetic time histories that were then used to develop site-specific response spectra (AMEC Earth and Environmental, 2003). Once the base isolators were designed, the design of shear walls was developed. Shear walls were placed vertically throughout the exterior and interior of the building (Fig. 6). The amount and length of the shear walls were greatly reduced by using base isolation and site-specific response spectra, and this allowed the architects greater latitude on placement of the shear walls within the building. Shear walls were located without disturbing the historic fabric of the building, and the integrity of the Capitol was preserved (Reaveley Engineers and Associates, 2003). The drum on top of the Capitol was also reinforced with shear walls, which were extensive because of the poor quality of the concrete (Fig. 6). New shear walls were placed on the interior of the drum wall, and steel rods were inserted from the shear walls through the existing drum wall and into a new exterior concrete wall. This design provided added strength to the drum and use of its existing concrete continued, reducing costs and schedule by eliminating the need to remove the old drum and construct a new one (Reaveley Engineers and Associates, 2003).

Ogden). The remainder of this route will travel across the Farmington Siding landslide complex for a better view. For a shorter, more direct route to stop 3, leave I-15 at exit 326 and travel north on U.S. 89, turning on Main Street as described below. If continuing on the longer route, take I-15 north 10 mi (16 km) to exit 325 (Lagoon Drive–Farmington), bear right to continue north on 200 West ~1.0 mi (1.6 km), then turn left (west) on State Street (U.S. 227). Follow State Street ~0.5 mi (0.8 km), crossing over the freeway, and bear right as State Street merges with 100 North. Continue west on 100 North ~1.0 mi (1.6 km); 100 North changes to Clark Lane. Turn right on 1525 West ~0.7 mi (1.1 km). Turn left on Burke Lane ~0.3 mi (0.5 km), continue to the right ~0.3 mi (0.5 km) as Burke Lane changes to 1875 West, and turn left on 900 North. Continue ~0.1 mi (0.2 km), turn right on 2000 West, continue ~0.3 mi (0.5 km) over the railroad crossing and past the pond on the right, and turn right on Shepard Lane (Utah 106). Travel east on Shepard Lane ~1.4 mi (2.3 km), crossing under I-15 and passing the Oakridge Country Club, and turn left on U.S. 89 ~0.9 mi (1.4 km). Turn right on Main Street (Utah 272) ~0.4 mi (0.6 km) and park on Main Street between Somerset Street and Leonard Lane.

Directions to Stop 3 From the State Capitol parking lot, turn right on Columbus Street and return to the intersection with 500 North. From the intersection, bear left on Victory Road for ~1.2 mi (1.9 km) until it merges with Beck Street. Continue northwest on Beck Street ~1.7 mi (2.7 km) to the entrance to I-15 North (toward

The prehistoric Farmington Siding landslide complex comprises some of the largest landslides triggered by earthquakes in the United States. The landslide complex covers an area of ~7.5 mi2 (19.5 km2) in southeastern Davis County, and is one of over a dozen late Pleistocene–Holocene features along the Wasatch Front that have been mapped as possible liquefaction-induced lateral spreads (Van Horn, 1975; Anderson et al., 1982; Nelson and Personius, 1993; Harty and Lowe, 2003). From the location of this field trip stop near the landslide main scarp, the landslide complex extends southwestward to Great Salt Lake. The Weber segment of the Wasatch fault zone lies ~0.5 mi (0.8 km) to the east. Data in this discussion are primarily from Hylland and Lowe (1998) and Harty and Lowe (2003).

Figure 6. Seismic design model for structural modifications to the Utah State Capitol, showing the location of shear walls (Reaveley Engineers and Associates, 2003). The shear walls (dark gray) reinforce the Capitol perimeter and interior, as well as the drum on top of the Capitol.

Stop 3—Liquefaction-Induced Farmington Siding Landslide Complex

Physical Setting and Geology The Farmington Siding landslide complex is in a gently sloping area underlain at shallow depths primarily by fine-grained, stratified, latest Pleistocene to Holocene lacustrine deposits of Lake Bonneville and Great Salt Lake. Ground slopes within the landslide complex range from ~0.4%–0.8% and adjacent to the complex range from ~1%–2% along the flanks to 6%–11% in the crown area. The deposits involved in landsliding consist of interbedded, laterally discontinuous layers of clayey to sandy silt, well-sorted fine sand to silty sand, and minor clay and gravel. The crown area is underlain by Lake Bonneville sand and silt deposits and is at an elevation of ~4400 ft (1340 m) in the vicinity of the city of Farmington. The toe may have been encountered beneath Great Salt Lake during a drilling project in Farmington Bay to

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test foundation conditions for a proposed water storage reservoir (Everitt, 1991). Soil grain size distribution, standard penetration resistance, and groundwater depth noted on logs of geotechnical boreholes (Anderson et al., 1982) indicate liquefiable deposits in the shallow subsurface beneath the landslide complex. These data, as well as data from three boreholes drilled in an attempt to correlate beds beneath and adjacent to the landslide complex (Miller et al., 1981), indicate a possible landslide failure zone within a depth range of 14–40 ft (4–12 m). This zone locally corresponds to the contact between a relatively dense transgressive lacustrine sequence consisting of nearshore sand and gravel deposited during the early part of the Bonneville paleolake cycle, or possibly pre-Bonneville alluvium, and overlying loose-soft, offshore, finegrained sediment subsequently deposited in deeper water. Geomorphic features within the landslide complex include scarps, hummocks, closed depressions, and transverse lineaments. Well-preserved lateral and main scarps in the northern part of the complex range from ~10–40 ft (3–12 m) high. Hummocks and closed depressions are present across most of the complex, but are more common in the northern part (Fig. 7). Hummocks in the northern part are morphologically distinct, having as much as ~20 ft (6 m) of relief and lateral dimensions locally exceeding 1000 ft (300 m). Hummocks in the southern part are morphologically subtle, generally having <~6 ft (2 m) of relief. Landslide Characteristics Trenches were excavated in several areas to examine the style of deformation and characterize the type of mass movement of the landslide complex (Hylland and Lowe, 1998; Harty and Lowe, 2003). Observed subsurface deformation of lacustrine deposits includes inclined strata, gentle to strong folding, highand low-angle shear surfaces, and loss of original bedding due to

Figure 7. Aerial view of hummocky landslide terrain on the northern part of the Farmington Siding landslide complex. View is to the northwest, with Interstate 15 near the middle of the picture.

liquefaction. Small sand dikes are present locally, some of which were injected along shear planes. Mass movement within the landslide complex was probably a combination of lateral spread and flow. By excavating trenches across hummock flanks and adjacent ground in the northern part of the complex, Harty and Lowe (2003) determined the hummocks are relatively intact “islands” of lacustrine strata surrounded by liquefied sand, which resulted from flow failure. Other evidence for flow failure includes the existence of a landslide main scarp up to 40 ft (12 m) high, overall negative relief in the head region of the complex indicating evacuation of a large volume of material, and overall positive relief in the distal region of the complex indicating accumulation of landslide material. Landslide Timing and its Relation to Surface-Faulting Earthquakes The landslide deposits can be grouped in two age categories relative to the age of the Gilbert shoreline complex of Lake Bonneville, which formed between 11,000 and 10,000 14C yr B.P. (Currey, 1990). The northern part of the landslide complex truncates the Gilbert shoreline (Van Horn, 1975), indicating major post-Gilbert movement. However, the Gilbert shoreline can be traced across the southern part of the landslide complex (Anderson et al., 1982; Harty and Lowe, 2003), indicating pre-Gilbert movement in this area. Relative timing information and radiocarbon soil ages indicate at least three, and possibly four, episodes of large-scale liquefaction-induced landsliding: the first sometime between 14,500 and 10,900 14C yr B.P., the second just prior to 7310 ± 60 14C yr B.P., the third (?) sometime prior to 5280 ± 60 14C yr B.P., and the fourth between 2340 ± 60 and 2440 ± 70 14C yr B.P. Hylland and Lowe (1998) considered the timing of these landslide events within the context of paleoclimatic and lacustral fluctuations and observed that landsliding was associated with climate-induced highstands of Great Salt Lake. The apparent correspondence between landslide events and lacustral highstands suggests that landsliding may have occurred under conditions of relatively high soil pore-water pressures, and possibly increased artesian pressures, associated with rising lake and groundwater levels. Many features (for example, evidence of lateral spread, flow failure of gentle slopes, sand dikes, deposits susceptible to liquefaction, and proximity to faults with recurrent Holocene activity) indicate landsliding was likely triggered by strong earthquake ground shaking. Comparison of the timing of surface-faulting earthquakes on the active segments of the Wasatch fault zone with the timing of Farmington Siding landslides indicates a close correspondence between landsliding and certain earthquakes (Fig. 8). Within uncertainty limits, surface-faulting earthquakes on the Brigham City segment coincide with all four possible landslide events. Surface-faulting earthquakes on the Weber, Salt Lake City, and Provo segments also coincide with the more recent landslide events. However, Hylland and Lowe (1998) and Hylland (1999) determined that large (surface-faulting) earthquakes on the nearby Weber segment of the Wasatch fault zone are significantly

Geologic hazards of the Wasatch Front, Utah more likely than earthquakes on other segments to trigger the widespread liquefaction-induced ground failure and large lateral displacements characteristic of the Farmington Siding landslide complex. The lack of evidence for liquefaction-induced landsliding triggered by the most recent surface-faulting earthquake on the Weber segment, which occurred during a relative lowstand of the lake, indicates that liquefaction-induced landsliding at the Farmington Siding landslide complex is tied not only to the likelihood of a surface-faulting earthquake on the Weber segment, but also to concurrent hydrogeologic conditions. Directions to Stop 4 Continue south on Main Street (Utah 272) ~0.4 mi (0.6 km), turn left on 1400 North ~0.5 mi (0.8 km), and turn right on Compton Road. Drive ~0.4 mi (0.6 km) on Compton Road and park. Stop 4—Compton Bench Fire-Related Debris Flows, Farmington

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Range had no protective structures. Due to the heightened debrisflow–flooding hazard, the U.S. Forest Service and U.S. Natural Resources Conservation Service undertook hazard-reduction measures to protect the watersheds. These protective measures, completed in late fall 2003, involved reseeding within the burn area and constructing sediment fences within channels below the small, unprotected drainage basins where houses were at risk. Physical Setting and Geology The small drainages that produced the Compton Bench debris flows lie along the foot of the Wasatch Range and rise in elevation from ~4600–4800 ft (1400–1500 m) at their mouths to 6400 ft (2000 m). The bedrock in these small drainages comprises the Precambrian Farmington Canyon Complex (Bryant, 1988), which includes schist, gneiss, and quartzite. Above 5200 ft (1600 m), the elevation of the Bonneville shoreline of Pleistocene Lake Bonneville, the slopes consist of colluvium with minor rock outcrops. Below 5200 ft (1600 m) lacustrine sand and gravel deposits of Lake Bonneville overlie the Farmington Canyon Complex (Nelson

Debris flows consist of rock, soil, and other debris that mix with water from intense thunderstorms or spring snowmelt and travel downslope at high speeds. Because of their considerable mass and speed, debris flows are a threat to life and can damage anything in their path, particularly buildings, roads, and utility connections. Debris flows are usually much more damaging than flash floods, but are generally more restricted in the area they impact. The Compton Bench fire-related debris flows are an excellent example of small-volume debris flows produced from small drainage basins. The lower slopes of the Wasatch Range above and east of Farmington City were burned in the Farmington fire of July 2003. Intense thunderstorms on 6 April 2004 triggered five small debris flows in the Compton Bench area of Farmington, a gently sloping area of coalesced alluvial fans. The Farmington Fire and Post-Fire Hazard Assessment The Compton Bench debris flows occurred eight months after a wildfire in the mountains above Farmington. The Farmington wildfire was a human-caused fire that burned ~2000 acres (800 ha) of U.S. Forest Service and private land on the lower western flank of the Wasatch Range above Farmington (U.S. Forest Service, 2003). The fire affected drainage basins above Farmington and Rudd Canyons and several smaller drainage basins. A debris-flow hazard assessment conducted by the UGS following the fire recognized a heightened debris-flow–flooding hazard for the tributaries within the burn area based on evidence of previous debris flows, erodible sediment stored in channels, burned hillslopes capable of generating rapid runoff, and rapid snowmelt and/or thunderstorm-rainfall potential (Giraud, 2003). Since Farmington and Rudd Canyons had previously produced damaging debris flows, debris basins had been constructed at the mouths of these canyons and provided protection for downslope development on alluvial fans. However, small drainage basins with areas of 5–80 acres (2–30 ha) along the lower flank of the Wasatch

Figure 8. Comparison of the timing of Farmington Siding landslides (shaded areas) with Wasatch fault zone surface-faulting earthquakes (top) and Great Salt Lake fluctuations (bottom) (modified from Hylland and Lowe, 1998). Dashed lines for landslide events indicate limiting ages; ages for landslides 2 and 3(?) are minimum limiting ages. Earthquake consensus preferred ages (dashed lines) and uncertainty limits (boxes) from Lund (2005); Great Salt Lake hydrograph after Murchison (1989). Age picks for calendar-calibrated time scale (top) were determined using the calibration program of Stuiver et al. (2005) and are approximate.

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and Personius, 1993). Alluvial fans formed at the mouths of drainages following the recession of Lake Bonneville. The Compton Bench Debris Flows The Compton Bench debris flows were triggered by intense thunderstorm rainfall between 8:30 and 9:00 p.m. on 6 April 2004. Two rain gauges about a mile from where the debris flows initiated recorded relatively small amounts of rainfall. A Davis County rain gauge at the Farmington Canyon debris basin at 4800 ft (1500 m) elevation recorded 0.17 in. (0.43 cm) in 17 min, and a U.S. Forest Service rain gauge in Rudd Canyon at 5200 ft (1600 m) elevation recorded 0.57 in. (1.4 cm) in 23 min. The debris flows initiated at elevations of ~6000–6200 ft (1800–1900 m), and the rainfall amounts and intensity were likely higher in the initiation areas than measured at the rain gauges. Rainfall measurements in other burned areas indicate that relatively small amounts of intense thunderstorm rainfall in the range of 0.27–0.35 in/h (0.7–0.9 cm/ h) are capable of triggering fire-related debris flows (McDonald and Giraud, 2002; Cannon et al., 2003). Other factors contributing to debris flows include steep slopes, ample supply of channel sediment, and increased runoff caused by soils with high moisture content from recent snowmelt. A traverse up the small drainage basins was conducted the morning after the debris flows to assess initiation processes and sediment bulking characteristics and to assist Farmington City in determining the potential for future debris flows. The hillslopes in the upper parts of the small drainage basins showed evidence indicating the debris flows began as intense runoff, and sheetwash erosion concentrated as rills and quickly flowed into the drainage basin channels. The debris flows evidently began entraining sediment in the upper parts of the drainage basin channels and continued to bulk sediment progressively downstream, through erosion and scour of the main channels. Below the Bonneville shoreline, abundant, loose, easily erodable sediment in the channels was bulked into the flows. At one locality, channel erosion threatened a section of a Weber Basin Water Conservancy District aqueduct running along the mountain front. Deposit volumes ranged from 200 to 1500 yds3 (150– 1100 m3). Debris-flow deposits include thin levees along alluvial fan channel margins and thin splays and lobes on small alluvial fans. The lateral deposit margins were flat rather than steep, indicating high water content for the flows. Two of the larger flows converged on Compton Road and flowed west across the road, covering parts of three lots with sediment (Fig. 9). These two flows traveled ~3000 ft (910 m) and dropped 1200 ft (370 m) in elevation. The flows traveled down shallow alluvial fan channels, collapsing the sediment fences in the channels. The high water content and highly fluid character of these flows promoted a longer runout. Sediment deposition was mostly restricted to streets and yards, but damage also occurred to several vehicles, garages, and homes (Fig. 10). Most of the deposits had a maximum thickness of 0.5 ft (0.2 m) unless the sediment flowed up against a building or other flow barrier, where sediment burial depths were up to 3 ft (0.9 m).

Compton Bench Risk-Reduction Structures A reconnaissance of the small drainage basins following the 6 April 2004 debris flows indicates that ample sediment is available for future debris flows. The potential for future debris flows prompted Farmington City, Davis County, and the U.S. Forest Service to collaborate and construct debris basins east of Compton Road to reduce the risk to lots impacted by the 6 April debris flows. Two small debris basins were constructed on the alluvial fan. An upper basin intercepts water and debris from two channels and connects to a lower basin where a high-level water outlet conveys water to the stormwater system. A berm downslope of the lower basin along the east side of Compton Road provides additional sediment storage if the capacity of the two debris basins is exceeded. Directions to Stop 5 Continue south on Compton Road ~0.4 mi (0.6 km), turn right on 1100 North, and follow 1100 North ~0.3 mi (0.5 km) as it changes to Quail Flight, which ends at Quail Circle. Turn right on Quail Circle and take a quick turn left on Quail Run Road at the end of Quail Circle. Take Quail Run Road ~0.1 mi (0.2 km), turn right on Main Street ~0.1 mi (0.2 km), and turn left on Shepard Lane. Continue 0.4 mi (0.6 km) on Shepard Lane, turn left (south) on U.S. 89 ~1.2 mi (1.9 km) and turn right (west) on I-15 northbound ~8.3 mi (13.4 km). Take Exit 332 (Syracuse) from I-15, turn left (west) on Antelope Drive (U.S. 108), and proceed west ~6.8 mi (10.9 km) to the Antelope Island State Park fee station following signs for Syracuse and Antelope Island. Cross the causeway to Antelope Island, ~6.8 mi (10.9 km), and bear left on the paved road. Follow the paved road ~1.2 mi (1.9 km) to an intersection and turn right. Continue on the road ~0.9 mi (1.4 km) to Buffalo Point, and park at the Bridger Bay Campground. Stop 5—Flood Hazards Related to Great Salt Lake A substantial portion of the Wasatch Front metropolitan area, from Salt Lake City north to Brigham City, is developed on a relatively narrow plain sandwiched between the eastern shore of Great Salt Lake and the Wasatch Range (Fig. 1). Low-lying coastal areas adjacent to Great Salt Lake’s 60-mi-long (100-km) eastern shore are subject to flooding from three main processes: (1) lake-level rise in response to multiyear wet cycles; (2) lake seiches; and (3) tectonic subsidence associated with surface faulting on the Wasatch fault zone. The most recent damaging flood took place between 1983 and 1987, when above-normal precipitation caused the lake level to rise to a new twentieth century high of 4212 ft (1284 m) above sea level in 1986 and again in 1987. Flooding associated with this highstand caused an estimated $240 million in damage to shoreline development. As a result of the recent drought (1999–2004), the lake level declined to a lowstand of ~4194 ft (1279 m) at the end of summer 2004. However, the winter of 2004–2005 produced above-normal precipitation and snowpack over most of the lake’s drainage basin, and the lake is rising again.

Geologic hazards of the Wasatch Front, Utah

Figure 9. Looking downslope at the three lots damaged by the 6 April 2004 debris flows. Crews are removing sediment from the lots and Compton Road. The debris flow exceeded the capacity of the small channel on the alluvial fan in the foreground and flowed onto the fan surface. The alluvial fan channel gradient decreases from 9° to 7° downfan and average channel depth is 2.5–3.0 ft (0.75–0.90 m).

Geologic Setting Great Salt Lake is the largest closed-basin lake in the Great Basin and the sixth largest lake in the United States; only the Great Lakes are larger. The complex structural basin occupied by Great Salt Lake was created during the episode of extensional tectonics, dating from the past 17 m.y. (Hintze, 1988), that gave rise to the Basin and Range Province. The lake occupies the lowest spot in a 22,000-mi2 (57,000-km2) drainage basin that extends beyond the northeastern Great Basin into the adjacent middle Rocky Mountains. Runoff from the Jordan, Weber, Ogden, and Bear Rivers is the major water input; evaporation from the lake surface is the major output from the system. Salinity of the water varies inversely with lake level from 5% to 28%. The present lake basin, and the deeper portions of its Pleistocene predecessor Lake Bonneville, coincide with three interconnecting fault-bounded grabens (Stokes, 1980). The intervening horst blocks give rise to the various islands in the lake, of which Antelope Island is the largest. Structurally, the island is the exposed tip of a tilted horst block, bounded on the west by the Antelope Island section of the Great Salt Lake fault zone (Willis et al., 2000). Lake Bonneville and Its Shorelines Lake Bonneville was the largest and most recent of the pluvial lakes that occupied the Bonneville Basin during the late Pleistocene (Fig. 11) (Currey, 1990; Oviatt et al., 1992; Oviatt, 1997). The elevations of the prominent shorelines of Lake Bonneville and Great Salt Lake are given in Table 3. Lake Bonneville began to rise from levels close to those of modern-day Great Salt Lake ca. 28,000 14C yr B.P. The lake rose gradually, but experi-

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Figure 10. Debris flows crossed Compton Road and buried the driveway, pushing the trucks and breaking the lower garage door panel. Water and sediment flowed into the garage, damaging the contents inside. Sediment burial depths are 1–3 ft (0.3–0.9 m).

Figure 11. Hydrograph of estimated lake levels in the Bonneville basin for the past 160,000 yr (from Hylland et al., 1997; originally modified from Currey and Oviatt, 1985). Ss—Stansbury shoreline, Bs—Bonneville shoreline, Ps—Provo shoreline, Gs—Gilbert shoreline.

enced a major, climatically induced oscillation between 21,000 and 20,000 14C yr B.P. that produced the Stansbury shoreline (Oviatt et al., 1990). The lake eventually resumed its rise, but the rise slowed as the lake level approached an external basin overflow threshold in southern Idaho. Lake Bonneville reached this threshold and occupied its highest shoreline, which Gilbert (1875) named the Bonneville beach, after 15,500 14C yr B.P. The lake remained at this level perhaps as late as 14,500 14C yr B.P., when catastrophic failure at Red Rock Pass, Idaho, an event known as the Bonneville Flood (O’Conner, 1993), lowered the lake level 350 ft (110 m) in less than two months, stabilizing at

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TABLE 3. ELEVATIONS AND AGES OF MAJOR SHORELINES AND LEVELS OF THE LAKE BONNEVILLE-GREAT SALT LAKE SYSTEM* Shoreline or lake level

Elevation (ft)

Date or age

Historic high Historic average Historic low Holocene high Holocene low Gilbert shoreline Provo shoreline Bonneville shoreline Stansbury shoreline

4212 4202 4191 4221 4180 (?) 4250 4740 5200 4500

1873, 1986, and 1987 N.A. 1963 ca. 3000 14C yr B.P. ca. 6000 14C yr B.P. ca. 11,000–10,000 14C yr B.P. ca. 14,500–14,000 14C yr B.P. ca. 15,500–14,500 14C yr B.P. ca. 21,000–20,000 14C yr B.P.

*Modified from Currey et al., 1984.

the newly established Provo shoreline. After 14,000 14C yr B.P., Lake Bonneville became a reduced, hydrologically closed system as overflow ceased when the climate warmed in response to a northward shift in the mean position of the westerly storm tracks following melting of the Laurentide and Cordilleran ice sheets (Oviatt, 1997). By ca. 12,000 14C yr B.P., the lake dropped to levels possibly at or lower than those of modern-day Great Salt Lake (Oviatt et al., 1992), but Lake Bonneville then transgressed to form the Gilbert shoreline between 11,000 and 10,000 14C yr B.P. Since the moderate rise to the Gilbert, possibly correlative with the climatic cooling of the Younger Dryas event (Oviatt, 1997), lake levels in the much-reduced Great Salt Lake have fluctuated more modestly during Holocene time. The Bonneville shoreline is a prominent geomorphic feature of the Wasatch Front. Easily visible from much of Salt Lake Valley, the elevation of the Bonneville shoreline varies from 5161–5216 ft (1573–1590 m) due to a combination of post-lake isostatic rebound and faulting (Hylland et al., 1997). Many shoreline features and other scientifically important “geoantiquities” related to Lake Bonneville have been covered, quarried, or otherwise lost due to population growth and associated urbanization along the Wasatch Front (Chan and Milligan, 1995). Historic Lake-Level Fluctuations and Flooding During the 1980s The following discussion of historic lake levels and flood events is abstracted from Arnow and Stephens (1990), Atwood et al. (1990), Atwood and Mabey (1995), Hylland et al. (1997), and Atwood (2002). The United States Geological Survey (USGS) has been monitoring water levels in Great Salt Lake since 1875. These data combined with Gilbert’s (1890) pregauge hydrograph (1847–1875) provide a nearly complete record of lake-level fluctuations since the time of Mormon settlement in the area. The historic average level of Great Salt Lake is ~4202 ft (1281 m); at this level, the lake covers an area of ~1700 mi2 (4400 km2). The 150-year-plus record of lake levels reveals long-term fluctuations that correlate directly with variations in the amount of precipitation within the lake’s drainage basin. Prior to the mid-

1980s, the historic high of Great Salt Lake was ~4212 ft (1284 m), which was reached in the early 1870s. Over the next 90 yr, the lake slowly dropped until reaching a historic low of 4191 ft (1277 m) in 1963, covering only 950 mi2 (2460 km2). September 1982 was the wettest month in the history of precipitation measurements at Salt Lake International Airport. This record wet month began a four-year wet cycle, in part related to strong El Niño events in 1982–1983 and 1986–1987 (Alder, 2002), over most of the Great Salt Lake drainage basin. At the beginning of September 1982, the lake was at an elevation of ~4200 ft (1280 m). From September 1982 to June 1983, the lake rose 5.2 ft (1.6 m), a record seasonal fluctuation. The lake continued to rise and reached a highstand of 4212 ft (1284 m) in 1986 and again in 1987, equaling or slightly exceeding the level reached in the 1870s. At this elevation the lake covered ~3300 mi2 (8547 km2), resulting in flooding of low-lying coastal areas. Post-flood mapping of the debris lines on Antelope Island created during this flood indicate significant wind setup and wave runup above the static level of the lake (Atwood, 1994; Atwood and Mabey, 1995; Atwood, 2002). On Antelope Island, the elevation of debris lines associated with the 1980s highstand of 4212 ft (1284 m) ranges from 4212–4218 ft (1284–1286 m). The magnitude and spatial variation in this process of shoreline superelevation needs to be considered when defining the lake’s flood plain (Atwood, 2002). From 1982 to 1986 the lake rose 12 ft (3.7 m), doubling its volume and flooding ~500,000 acres (202,500 ha) of the lake’s historic flood plain. The flooding and associated wave erosion damaged or destroyed public and private resources and facilities, including damage to the Rose Park industrial area, Interstate 80, the Antelope Island causeway, park pavilions at Antelope Island State Park (Fig. 12), and salt company evaporation ponds and dikes. The overall economy of the State of Utah was dramatically impacted, with total flood-related losses estimated at $240 million (Austin, 2002). In response to the unprecedented flood damage, the State of Utah employed two measures in an attempt to quickly lower the level of Great Salt Lake: (1) breaching the Southern Pacific Railroad (SPRR) causeway in 1984, and (2) pumping lake water into the West Desert in 1987 (Gwynn, 2002). By 1983, the SPRR causeway had created a 3.5-ft (1.1-m) lake-level differential between the north (lower) and south (higher) arms of the lake due to the impermeable nature of the causeway and the abnormally high inflows of water into the south arm of the lake during the early 1980s. The causeway was breached on 1 August 1984 and within several months the lake level differential had been reduced to ~1 ft (0.3 m). However, the south arm of the lake continued to rise. In 1986, the State of Utah decided to pump lake water westward into the Great Salt Lake desert, adjacent to the Bonneville Salt Flats. The West Desert Pumping Project was an attempt to increase evaporation rates and remove 690,000 acre-ft (84,870 ha-m) from the lake. Three large pumps were installed near the south end of the Hogup Mountains that lifted water from the lake into a 4.1-mi (6.6-km) canal and transported it westward into a shallow topographic depression known as the West Pond. By the

Geologic hazards of the Wasatch Front, Utah end of 1988, the lake level had dropped ~5.4 ft (1.6 m) due to a combination of pumping, evaporation, and decreased inflow resulting from two drier-than-average years (Hylland et al., 1997). Pumping (April 1987 to June 1989) probably lowered the level of the lake ~2 ft (0.6 m), or 36% of the total decline during that period (Gwynn, 2002). The pumping facility is maintained in a state of ready reserve for future high lake levels. Lake Seiches A seiche is a free or standing-wave oscillation of the water level in an enclosed or semi-enclosed basin, such as a lake, bay, or harbor. Wind-driven lake setup, landslides, earthquake-induced ground shaking, or surface faulting on the lake floor may create a seiche. The large waves created by surface faulting beneath the body of water are sometimes called surges to differentiate them from the smaller oscillations caused by wind setup or ground shaking (Myers and Hamilton, 1964). Wind seiches in Great Salt Lake are a fairly common occurrence, are well documented by USGS lake-level monitoring (Atwood, 2002), and have been studied in some detail (Wang, 1978). Southerly winds in advance of a passing cold front will induce lake setup in the north and lake set-down along the southern shoreline. When the wind velocity drops, the water level will oscillate, with a fundamental period of ~6 h, as the lake tries to regain equilibrium (Fig. 13). Wang (1978) found that significant lake setup requires a wind velocity of 10 knots (18.5 km/h) for 12 h duration. However, Atwood (2002) suggests that wind setup in the southern arm of Great Salt Lake may not require steady, strong winds of long duration. The maximum amplitude of winddriven seiching in Great Salt Lake is ~2 ft (0.6 m) (Wang, 1978). However, this could cause significant wave damage if superimposed on an already high lake level. The potential for earthquake-induced seiching in Great Salt Lake is poorly understood, but Lowe (1993) makes a strong case,

Figure 12. Debris from picnic facilities at north end of Antelope Island destroyed during mid-1980s rise of Great Salt Lake.

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based on historical lake-level data and earthquake accounts, that the 1909 Hansel Valley earthquake (M 6+) generated a 12-ft (3.7-m) wave that overtopped the Lucin Cutoff railroad trestle. Pechmann (1987) notes the exceptional size of this wave and that the larger 1934 M 6.6 Hansel Valley earthquake did not generate seiching in Great Salt Lake. He concludes that the epicenter of the 1909 earthquake may have actually been beneath Great Salt Lake and the reported wave was in fact a surge related to ground rupture beneath the lake, as opposed to ground shaking from a more distant earthquake. A future surge of similar size along the southern or eastern shore of Great Salt Lake would be a significant and potentially damaging event. However, for a scenario M 7 earthquake along the Salt Lake City segment of the Wasatch fault zone, neither seiching related to ground shaking nor a surge related to surface faulting along the Salt Lake City segment would be significant (Solomon et al., 2004). Tectonic Subsidence Several papers and reports address potential lake margin flooding caused by earthquakes (e.g., Lowe, 1993). Two major studies on the potential impacts of earthquake-induced tectonic deformation, tilting, and flooding along the Wasatch fault zone were published by Keaton (1986) and Chang and Smith (1998).

Figure 13. Hydrographs for Promontory Point (north) and Silver Sands (south), and wind speed at Salt Lake International Airport, showing wind seiches on Great Salt Lake over a five-day period (from Hylland et al., 1997; originally modified from Lin and Wang, 1978).

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Chang and Smith (1998) refined previously published deformation models by inputting more accurate elevation data into an elastic, three-dimensional boundary-element model to simulate ground-surface deformation associated with large, surface-faulting earthquakes on the Weber and Salt Lake City segments of the Wasatch fault zone. The study by Chang and Smith (1998) demonstrates that flooding from Great Salt Lake resulting from tectonic subsidence or ground tilt associated with a large surfacefaulting earthquake on the Wasatch fault zone could be a significant hazard for developed areas of Salt Lake and Davis Counties. The flooding hazard associated with such an earthquake is based on the premise that the topography of low-lying areas along the eastern shore of Great Salt Lake can be permanently altered by subsidence of the hanging wall, resulting in an eastward migration of the lake shoreline and inundation of adjacent areas. Superimposing the effects of any accompanying surges and

seiches would cause even greater inland flooding (Atwood and Mabey, 1995). A secondary hazard would be the adverse effects of backtilting on the many gravity-driven water and sewer lines that cross the fault zone. The worst-case scenario modeled by Chang and Smith (1998) (based on a M 7.5 earthquake on the Salt Lake City and Weber segments of the Wasatch fault zone similar to the 1959 Hebgen Lake, Montana, earthquake and Great Salt Lake at a level equal to its historic highstand of 4212 ft [1284 m]) results in a 3.5-mi (5.6-km) southeastward displacement of the shoreline onto land containing commercial developments and major transportation corridors. The degree of flooding that results from their models is very sensitive to the level of Great Salt Lake at the time of the earthquake. Recent GIS-based mapping (Hernandez, 2004) of the potential extent of flooding in southern Davis County demonstrates the vulnerability of the I-15 corridor in that area (Fig. 14). Chang and Smith (1998) state their results should be used with caution because of the lack of site-specific data in their models. In addition, Keaton (1986) indicates that the amount of hanging wall deformation produced by the 1959 Hebgen Lake earthquake (~20 ft [6.1 m] of vertical displacement) may be more than twice what would be expected for a slightly lower magnitude earthquake on either the Weber (Fig. 14) or Salt Lake City (Fig. 4) segments of the Wasatch fault zone (~6.9 ft [2.1 m], based on paleoseismic studies). Recent discussions of the seismic records of the Hebgen Lake earthquake are also raising questions about the number and magnitude of seismic events that actually occurred on that day (J.R. Keaton, 2004, personal commun.). Therefore, the extent of tectonic-related lake-margin flooding as shown in Figure 14 may be overestimated. Directions to Stop 6 Retrace the route to I-15, turn right (S) on I-15, travel for ~19 mi (31 km) to Exit 316, and turn right (W) to I-215. Take I-215 west and south ~9.0 mi (14 km) to exit 20A, turn right (W) on State Route 201, travel for ~7.5 mi (12 km), and turn left on 8400 West (U.S. 111) in Magna. Continue south on 8400 West for ~5.0 mi (8.0 km), turn right (W) on 5400 South, and park after ~0.5 mi (0.8 km), before the road turns left into the gravel pit. Stop 6—Faulted Pleistocene (?) Deposits in Western Salt Lake Valley

Figure 14. Tectonic-related flood distribution based on the historic high stand (4212 ft [1284 m]) of Great Salt Lake and hypothesized ground deformation on the southern portion of the Weber segment of the Wasatch fault zone. The mapping assumes the worst-case scenario (the observed hanging wall deformation associated with the 1959 MS 7.5 Hebgen Lake, Montana, earthquake) proposed by Chang and Smith (1998) (modified from Hernandez, 2004).

The Wasatch fault zone has been extensively studied, exhibits evidence of repeated movement during the Holocene, and its hazards are relatively well known. However, other Quaternary faults are also found in and near the central Wasatch Front. Some have not moved during the past 10,000 yr, the conventional criterion to define an active fault, but Pleistocene faults with long recurrence intervals may pose significant seismic hazards as well. DePolo and Slemmons (1998) suggest that a longer time period, 130,000 yr, is a more appropriate criterion because most earthquake recurrence intervals in the Basin and Range

Geologic hazards of the Wasatch Front, Utah Province exceed 10,000 yr; the larger time period encompasses most recurrence intervals for faults in the province, and at least 50% of historical earthquakes in the province of M 6.5 or greater involved fault traces that lacked prior Holocene faulting. At Stop 6 we examine a possible Pleistocene fault exposure near the west margin of Salt Lake Valley that has no surface expression but may be considered active using the criteria of DePolo and Slemmons (1998). Possible Pleistocene Faulting in Western Salt Lake Valley Recent detailed geologic mapping in western Salt Lake Valley revealed faulted deposits of possible Pleistocene age on the east side of the Oquirrh Mountains (Biek et al., 2004) (Fig. 4). These faults include the Harkers fault, extending at least 6 mi (10 km) from the range front to the range interior and separating the Permian and Pennsylvanian Oquirrh Group in the footwall from Miocene (?) to Pleistocene alluvial fan deposits in the hanging wall. Other small faults, exposed at Stop 6, separate the alluvial fan deposits from tuff in the Jordan Narrows unit of the Salt Lake Formation near the northern end of the Harkers fault, 1.0 mi (1.6 km) east of the range front. Bryant et al. (1989) reported a fission-track age of 4.4 ± 1.0 Ma for a rhyolitic tuff in the Jordan Narrows unit, which, with its stratigraphic position above rocks of known Miocene age, indicates a Miocene to Pliocene age for the unit (Biek et al., 2004). Biek et al. (2004) differentiated the pre-Bonneville alluvial fan deposits in western Salt Lake Valley into two units based on morphology and paleosol development: (1) older alluvial fan deposits form relatively steep, deeply dissected, erosionally resistant remnants commonly overlain by a stage IV calcic paleosol, indicating an early to middle Pleistocene age for the upper part of the unit; and (2) younger pre-Bonneville alluvial fan deposits form relatively gently sloping aprons on piedmont slopes, are commonly truncated by the Bonneville shoreline, and are overlain by stage II or III calcic paleosols, suggesting a middle to late Pleistocene age. No fossil evidence has been found in the pre-Bonneville alluvial fan deposits, but glass shard analyses of tuff samples from the adjacent Copperton quadrangle (Biek et al., 2004) suggest a chemical correlation to the 6.4 ± 0.1 Ma Walcott Tuff, indicating that the oldest beds of the deposits may be as old as late Miocene. Faults Exposed at Stop 6 Faults are exposed in pre-Bonneville beds in a railroad cut ~3 mi (5 km) south of Magna (Fig. 15). Two of the faults juxtapose tuffaceous lake beds and pre-Bonneville alluvial fan deposits, and several smaller faults occur within each unit. The amount of displacement on the larger faults cannot be determined, but is at least ~20 ft (6 m), the height of the exposed cut slope. The tuffaceous lake beds are part of the Miocene to Pliocene Jordan Narrows unit of the Salt Lake Formation. The alluvial fan deposits are assumed to be of Miocene (?) to middle Pleistocene age because of their proximity to outcrops of that unit, but paleosol preservation is poor at the exposure. The cut

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slope is ~0.8 mi (1.3 km) northeast of the Lake Bonneville highstand shoreline and is ~120 ft (37 m) lower in elevation. A thin veneer of Bonneville gravel that is not displaced caps the exposure. Because of the uncertainty in the age of the alluvial fan deposits, the faults may have been active during the Pleistocene prior to the Lake Bonneville transgression, but displacement may be as old as Miocene. The bounding faults trend N10–20°E and form a small graben that has displaced the alluvial fan deposits down with respect to the tuffaceous lake beds. The exposed graben is ~400 ft (120 m) wide. Although the alluvial fan deposits are downdropped, their greater relative erosional resistance resulted in a topographic high, but no linear fault trace is evident on the ground surface. Lake Bonneville gravel accumulated on the high to form a curved spit ~2000 ft (600 m) long on the fan remnant. The spit helped to shelter an adjacent small lagoon ~400 ft (120 m) in diameter, whose fine-grained deposits underlie a small, shallow, grass-covered depression. Directions to Stop 7 Return to 5400 South and travel east ~7.4 mi (11.9 km), passing under I-215, to Redwood Road. Turn right on Redwood Road, go ~0.5 mi (0.8 km) to I-215, and take I-215 east ~7.0 mi (11.3 km) to Exit 6 (6200 South). Turn right onto 6200 South, proceed uphill, and continue south on Wasatch Blvd for ~4.0 mi (6.4 km). Bear right at the traffic signal on Wasatch Blvd as it diverges from Little Cottonwood Road, and continue ~1.1 mi (1.8 km). Turn right on 9800 South and then immediately turn right into the parking lot for G.K. Gilbert Geologic Interpretive Park.

Figure 15. Miocene to Pliocene tuffaceous beds of the Salt Lake Formation, Jordan Narrows unit, in the footwall of a normal fault with possible Pleistocene movement (Stop 6). Miocene (?) to middle Pleistocene alluvial fan deposits are in the hanging wall. The height of the exposure is ~20 ft (6 m).

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Stop 7—Surface Faulting on the Salt Lake City Segment of the Wasatch Fault Zone The Wasatch fault zone is one of the longest and most active normal-slip faults in the world. The Salt Lake City segment of the Wasatch fault zone trends through the densely populated Salt Lake Valley and poses a significant seismic risk to the Salt Lake City metropolitan area (Fig. 16). The segment extends for ~29 mi (46 km) from the Traverse Mountains on the south to the Salt Lake salient on the north. The Salt Lake City segment displays abundant geologic and geomorphic evidence for multiple surface-faulting earthquakes during Holocene time. Personius and Scott (1992) subdivided the Salt Lake City segment into three subsegments (from north to south): (1) the Warm Springs fault, (2) the East Bench fault, and (3) the Cottonwood section. Urbanization within Salt Lake Valley has obscured or modified scarps along the Salt Lake City segment in many places. Here at Stop 7, the Wasatch fault principally cuts prominent Pinedale-age glacial moraines of Little Cottonwood and Bells Canyons, the only areas in Salt Lake Valley where the glaciers reached the mountain front. The highest shoreline of Lake Bonn-

Figure 16. Salt Lake City segment of the Wasatch fault zone and locations of the Little Cottonwood Canyon (LCC), South Fork Dry Creek (SFDC), and Dry Gulch (DG) trench sites (from Black et al., 1996). Arrows indicate approximate segment boundaries.

eville is just to the west, and recent, unpublished work by Elliot Lips of the University of Utah indicates that the maximum glacial advance and occupation of the Bonneville shoreline were contemporaneous (E. Lips, 2005, personal commun.). Scarps and the graben at this locality thus represent surface faulting over about the past 16,000 yr. Little regard was given to surface-faulting hazards in Salt Lake County land-use planning prior to 1970, when an informal fault investigation and review process was implemented for new buildings. In 1989, Salt Lake County enacted the Natural Hazards Ordinance, now revised and renamed the Geologic Hazards Ordinance, and requires site-specific investigations to locate active faults and establish appropriate building setbacks prior to development (Batatian, 2002). At Stop 7, we will discuss surface-faulting hazards along the Salt Lake City segment and paleoseismic studies to evaluate them and will observe significant geologic and geomorphic features of the segment. Regional Setting of the Wasatch Fault Zone The Wasatch fault zone extends for 213 mi (343 km) from southeastern Idaho to north-central Utah (Machette et al., 1992). The fault zone generally trends north-south and, at the surface, can form a zone of deformation up to several hundred feet wide containing many subparallel west-dipping main faults and eastdipping antithetic faults. Schwartz and Coppersmith (1984) originally divided the Wasatch fault zone into six independent, seismogenic segments based on scarp morphology, surface-faulting patterns, range-crest morphology, geophysical evidence, and limited trenching information. Based on additional detailed trenching studies and geologic mapping, Machette et al. (1992) proposed a revised segmentation scheme consisting of 10 segments. The central five segments of the fault zone (Brigham City, Weber, Salt Lake City, Provo, and Nephi), along the western base of the Wasatch Range, each show evidence of two or more surface-faulting earthquakes in the past 6000 yr (Black et al., 2003). The Wasatch Range is a major north-south–trending mountain range in Utah, and the Wasatch fault zone forms a prominent west-facing escarpment along its base. The fault zone is the easternmost feature of the Basin and Range Province, which is characterized by a series of generally north-trending elongate mountain ranges, separated by predominately alluvial and lacustrine sediment-filled valleys and typically bounded on one or both sides by major normal faults (Stewart, 1978). Late Cenozoic normal faulting, a characteristic of the Basin and Range Province, began between ca. 17 and 10 Ma in the Nevada (Stewart, 1978) and Utah (Anderson, 1989) portions of the province. The faulting is a result of a roughly east-west–directed, regional extensional stress regime that continues to the present (Zoback, 1989). The Wasatch fault zone is also near the center of the Intermountain seismic belt, a generally north-south–trending zone of historical seismicity extending from northern Arizona to northwestern Montana (Smith and Sbar, 1974). At least 16 earthquakes of M 6.0 or greater have occurred within the Intermountain seismic belt since 1850, but none of them occurred

Geologic hazards of the Wasatch Front, Utah along the Wasatch fault zone (Arabasz et al., 1992). The largest of these earthquakes was the MS 7.5 event in 1959 near Hebgen Lake, Montana. Paleoseismic Studies and Earthquake History The Holocene chronology of surface-faulting earthquakes on the Salt Lake City segment has been the subject of paleoseismic studies for more than two decades (Swan et al., 1981; Lund and Schwartz, 1987; Schwartz and Lund, 1988; Lund, 1992). The earthquake chronology was initially based on paleoseismic investigations at Little Cottonwood Canyon in 1979 and South Fork Dry Creek in 1985 (Fig. 16). A subsequent study at Dry Gulch in 1991 discovered a previously unrecognized event on a scarp not trenched at the South Fork Dry Creek site, which suggested the surface-faulting chronology was incomplete. Black et al. (1996) conducted additional trenching at the South Fork Dry Creek site in 1994 to establish a complete surface-faulting earthquake chronology from at least the middle Holocene for the Salt Lake City segment. McCalpin (2002) excavated a megatrench at the original Little Cottonwood Canyon site in 1999 in an attempt to capture the entire post-Bonneville record of paleoearthquakes on the Salt Lake City segment. In 2003, the UGS, with funding from the National Earthquake Hazards Reduction Program, convened the Utah Quaternary Fault Parameters Working Group, a panel of expert paleoseismologists and seismologists, to make a comprehensive evaluation of the paleoseismic trenching data available for Utah’s Quaternary faults and, where the data permitted, assign consensus-preferred recurrence-interval and vertical slip-rate

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estimates for the faults and fault sections under review (Lund, 2005). As part of the review, the Working Group considered the paleoseismic trenching data available for the Salt Lake City segment. Earthquake timing. The paleoseismic trenching data include stratigraphic evidence for seven paleoearthquakes (events T through Z) younger than the Bonneville flood, ca. 17.2 ka, and possibly an eighth event (event S) that occurred while Lake Bonneville was at or near its highstand at the Bonneville shoreline. The Working Group’s consensus for the timing of surface faulting on the Salt Lake City segment is shown in Table 4. Timing for earthquakes W, X, Y, and Z is from Black et al. (1996) and comes from the South Fork Dry Creek and Dry Gulch trench sites. The confidence limits for each earthquake were increased to accommodate the full range of limiting 14C ages used to constrain the timing of the earthquakes (Lund, 2005). The Working Group believes that the resulting ranges account for both the laboratory and geologic uncertainty associated with the timing of each earthquake. McCalpin (2002) identified earthquakes S, T, U, and V at Little Cottonwood Canyon on the basis of a retrodeformation analysis of stratigraphic and structural relations exposed in the megatrench. No direct evidence (such as colluvial wedges, tectonic crack fills, or fault terminations) was found to document these earthquakes, and consequently their timing is only broadly constrained. Surface-faulting recurrence. Table 4 also shows the interevent recurrence intervals for the Salt Lake City segment, with associated confidence limits, determined from the earthquake chro-

TABLE 4. TIMING AND RECURRENCE INTERVALS OF HOLOCENE AND LATEST PLEISTOCENE SURFACE-FAULTING EARTHQUAKES ON THE SALT LAKE CITY SEGMENT OF THE WASATCH FAULT ZONE Surface-faulting event Event Z Event Y Event X Event W Event V

Event U Event T Event S (?)

Event timing* (cal yr B.P.)

Inter-event recurrence interval† (cal yr)

Post-event elapsed time (cal yr)

N.A. 1300 ± 650 N.A. 2450 ± 550 N.A. 3950 ± 550 N.A. 5300 ± 750 ca. 7.5 ka (after 8.8–9.1 ka but before 5.1–5.3 ka) ca. 9 ka (shortly after 9.5–9.9 ka) ca. 17 ka ca. 17–20 ka

N.A. N.A. 1150 ± 900 N.A. 1500 ± 800 N.A. 1350 ± 900 N.A. N.A.

1300 ± 650 N.A. N.A. N.A. N.A. N.A. N.A. N.A. N.A.

N.A.

N.A.

N.A. N.A.

N.A. N.A.

Note: Based on Lund (2005). *Timing for events W, X, Y, and Z from Black et al. (1996); timing for events S, T, U, and V from McCalpin (2002). † Confidence limits of recurrence intervals equal the square root of the sum of the squares of the individual confidence limits for each bracketing earthquake, rounded to the nearest 100 yr.

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nology. The weighted mean recurrence for the three most recent inter-event intervals (from event W through event Z) is 1300 ± 400 cal yr (Lund, 2005). The timing of earthquakes U and V on the Salt Lake City segment is broadly constrained. McCalpin (2002) reports the U–V and V–W interevent intervals are both ~2 k.y., resulting in a similarly broadly constrained mean recurrence for surface faulting of 2 k.y. for mid- to early Holocene time. The timing of event T is likewise broadly constrained. McCalpin (2002) reported a range in the interevent interval between events T and U of 7.1–9.6 k.y., with a mean of 8.4 k.y., indicating a long period of surface-faulting quiescence during earliest Holocene and latest Pleistocene time. However, McCalpin (2002) noted that the physical evidence of additional earthquakes in the gap may have been removed by alluvial fan erosion in the interval 9–10 ka and subsequently lost from the stratigraphic record. Based on currently available information on earthquake timing and variability in the length of individual interevent intervals, the Working Group’s preferred recurrence-interval estimate for the Salt Lake City segment is 1300 yr, with a minimum estimate of 500 yr and a maximum of 2400 yr (Lund, 2005). Vertical slip rate. Trenching investigations at Little Cottonwood Canyon and at the South Fork Dry Creek and Dry Gulch sites did not produce well-constrained net vertical displacement data. Swan et al. (1981) reported 47.6 (+30/−10) ft (14.5 [+10/ −3] m) of net vertical displacement across the Wasatch fault zone determined from a scarp profile measured along the crest of the Bells Canyon glacial moraine a few hundred meters south of Little Cottonwood Canyon. Scott (1988) reports the age of the moraine as 18–26 ka. The resulting slip rate using that age is a preferred rate of 0.03 in/yr (0.7 mm/yr), with a minimum estimate of 0.02 in/yr (0.4 mm/yr) and a maximum of 0.06 in/yr (1.4 mm/yr) (Lund, 2005). The Bells Canyon vertical slip rate is a long-term rate extending from the latest Pleistocene. Well-constrained vertical slip-rate data are lacking elsewhere on the Salt Lake City segment. Because the Bells Canyon long-term slip-rate estimate includes a possible period of seismic quiescence in the late Pleistocene (McCalpin, 2002), the Working Group’s slip-rate estimate for the Salt Lake City segment in the Holocene is higher than the longer-term rate at Bells Canyon. Based on currently available information on earthquake timing and displacement, the Working Group’s preferred vertical slip-rate estimate and confidence limits for the Salt Lake City segment in the Holocene is 0.05 in/yr (1.2 mm/yr), with a minimum estimate of 0.02 in/yr (0.6 mm/yr) and a maximum estimate of 0.16 in/yr (4.0 mm/yr) (Lund, 2005). SUMMARY This field trip exposes participants to examples of the wide variety of geologic hazards threatening the Wasatch Front urban corridor. Some of these hazards, such as landsliding examined at Stop 1 and debris flows at Stop 4, are geologically frequent—residents are vulnerable to similar hazards somewhere in the region

almost every year, with risks increasing during periods of aboveaverage precipitation, rapid spring snowmelt, and wildfires and cloudburst storms. Other hazards, such as the earthquake-related hazards of ground shaking discussed at Stop 2, liquefaction at Stop 3, and surface fault rupture at Stop 7, have longer recurrence intervals of several hundred years. Although infrequent, earthquake hazards can result in catastrophic consequences. One attempt to reduce earthquake hazards is the Utah State Capitol seismic retrofit (Stop 2). Flooding from Great Salt Lake (Stop 5) may be either climatically induced following cycles of several years, or earthquake induced, paralleling the pattern of Wasatch Front seismicity. Although hazards from large earthquakes are largely controlled by activity along the Wasatch fault zone, faults exposed at Stop 6 illustrate the potential for seismic hazards posed by faults with longer recurrence intervals. Whatever the cause of geologic hazards, residents of the Wasatch Front accept risk regardless of location. Our field trip, and related studies, is part of the effort to enhance awareness of the hazards. Industry and local governments can then use the results of these studies to implement hazard-reduction measures. REFERENCES CITED AGRA Earth and Environmental, 2000, Material testing services—existing material conditions: Salt Lake City, Utah, unpublished consultant’s report. Alder, W., 2002, The National Weather Service, weather across Utah in the 1980s, and its effect on Great Salt Lake, in Gwynn, J.W., ed., Great Salt Lake—an overview of change: Salt Lake City, Utah Department of Natural Resources Special Publication, p. 295–301. AMEC Earth and Environmental, 2003, Ground motion for design of seismic renovation, Utah State Capitol Building, Salt Lake City, Utah: Salt Lake City, Utah, unpublished consultant’s report. Anderson, L.R., Keaton, J.R., Aubrey, K., and Ellis, S.J., 1982, Liquefaction potential map for Davis County, Utah: Logan, Utah State University Department of Civil and Environmental Engineering and Dames & Moore Consulting Engineers, final technical report for the U.S. Geological Survey: Utah Geological Survey Contract Report 94-7, Utah Geological Survey Open-File Report 433, 50 p. Anderson, R.E., 1989, Tectonic evolution of the intermontane system—Basin and Range, Colorado Plateau, and High Lava Plains, in Pakiser, L.C., and Mooney, W.D., eds., Geophysical framework of the continental United States: Geological Society of America Memoir 172, p. 163–176. Arabasz, W.J., Pechmann, J.C., and Brown, E.D., 1992, Observational seismology and evaluation of earthquake hazards and risk in the Wasatch Front area, Utah, in Gori, P.L., and Hays, W.W., eds., Assessment of regional earthquake hazards and risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1500-D, p. D1–D36. Arnow, T., and Stephens, D., 1990, Hydrological characteristics of the Great Salt Lake, Utah 1847–1986: U.S. Geological Survey Water-Supply Paper 2332, 32 p. Ashland, F.X., 2002, Recurrent movement of the East Capitol Boulevard–City Creek landslide, Salt Lake City, Utah [abs.]: Association of Engineering Geologists 45th Annual Meeting Program with Abstracts, p. 54. Ashland, F.X., 2003, Characteristics, causes, and implications of the 1998 Wasatch Front landslides, Utah: Utah Geological Survey Special Study 105, 49 p. Ashland, F.X., Giraud, R.E., and McDonald, G.N., 2005, Groundwater-level fluctuations in Wasatch Front landslides and adjacent slopes, northern Utah: Utah Geological Survey Open-File Report 448, 22 p. Ashland, F.X., and McDonald, G.N., 2003, Interim map showing shear-wavevelocity characteristics of engineering geologic units in the Salt Lake City, Utah metropolitan area: Utah Geological Survey Open-File Report 424, 43 p. pamphlet, scale 1:75,000, CD-ROM. Atwood, G., 1994, Geomorphology applied to flooding problems of closedbasin lakes, specifically Great Salt Lake, Utah, in Morisawa, M., ed., Geo-

Geologic hazards of the Wasatch Front, Utah morphology and natural hazards—Proceedings of the 25th Binghamton symposium in geomorphology: Geomorphology, v. 10, p. 197–219. Atwood, G., 2002, Storm-related flooding hazards, coastal processes, and shoreline evidence of Great Salt Lake, in Gwynn, J.W., ed., Great Salt Lake—an overview of change: Salt Lake City, Utah Department of Natural Resources Special Publication, p. 43–53. Atwood, G., and Mabey, D.R., 1995, Flooding hazards associated with Great Salt Lake, in Lund, W.R., ed., Environmental and engineering geology of the Wasatch Front region: Utah Geological Association Publication 24, p. 483–493. Atwood, G., Mabey, D.R., and Lund, W.R., 1990, The Great Salt Lake—a hazardous neighbor, in Lund, W.R., ed., Engineering geology of the Salt Lake City metropolitan area, Utah: Utah Geological and Mineral Survey Bulletin 126, p. 54–58. Austin, L.H., 2002, Problems and management alternatives related to the selection and construction of the West Desert Pumping Project, in Gwynn, J.W., ed., Great Salt Lake—an overview of change: Salt Lake City, Utah Department of Natural Resources Special Publication, p. 303–312. Batatian, D., 2002, Minimum standards for surface fault rupture hazard studies: Salt Lake City, Salt Lake County Geologic Hazards Ordinance, Chapter 19.75, Appendix A, 11 p. Biek, R.F., Solomon, B.J., Keith, J.D., and Smith, T.W., 2004, Interim geologic maps of the Copperton, Magna, and Tickville Spring quadrangles, Salt Lake and Utah Counties, Utah: Utah Geological Survey Open-File Report 434, scale 1:24,000. Black, B.D., Hecker, S., Hylland, M.D., Christenson, G.E., and McDonald, G.N., 2003, Quaternary fault and fold database and map of Utah: Utah Geological Survey Map 193DM, scale 1:500,000, CD-ROM. Black, B.D., Lund, W.R., Schwartz, D.P., Gill, H.E., and Mayes, B.H., 1996, Paleoseismology of Utah, Volume 7—Paleoseismic investigation on the Salt Lake City segment of the Wasatch fault zone at the South Fork Dry Creek and Dry Gulch sites, Salt Lake County, Utah: Utah Geological Survey Special Study 92, 22 p. Borcherdt, R.D., 1994, Estimates of site-dependent response spectra for design - methodology and justification: Earthquake Spectra, v. 10, p. 617–653, doi: 10.1193/1.1585791. Bryant, B., 1988, Geology of the Farmington Canyon complex, Wasatch Mountains, Utah: U.S. Geological Survey Professional Paper 1476, 54 p., map scale 1:50,000. Bryant, B., Naeser, C.W., Marvin, R.F., and Mehnert, H.H., 1989, Ages of late Paleogene and Neogene tuffs and the beginning of rapid regional extension, eastern boundary of the Basin and Range Province near Salt Lake City, Utah: U.S. Geological Survey Bulletin 1787-K, 12 p., 1 plate, scale ~1:500,000. Cannon, S.H., Gartner, J.E., Holland-Sears, A., Thurston, B.M., and Gleason, J.A., 2003, Debris flow response of basins burned by the 2002 Coal Seam and Missionary Ridge fires, Colorado, in Boyer, B.D., Santi, P.M., and Rodgers, W.P., eds., Engineering geology in Colorado—contributions, trends, and case histories: Colorado Geological Survey Special Publication 55, 30 p. Chan, M.A., and Milligan, M.R., 1995, Gilbert’s vanishing deltas—a century of change in Pleistocene deposits of northern Utah, in Lund, W.R., ed., Environmental and engineering geology of the Wasatch Front region: Utah Geological Association Publication 24, p. 521–532. Chang, W.L., and Smith, R.B., 1998, Potential for tectonically induced tilting and flooding by the Great Salt Lake, Utah, from large earthquakes on the Wasatch Fault, in Lund, W.R., ed., Proceedings volume—Basin and Range Province Seismic-Hazards Summit: Utah Geological Survey Miscellaneous Publication 98-2, p. 128–138. Cooper Roberts Simonson Architecture, 2000, Utah State Capitol—Building and grounds restoration master plan and historic structures report: Salt Lake City, Utah, unpublished consultant’s report, two volumes, variously paginated. Currey, D.R., 1990, Quaternary paleolakes in the evolution of semidesert basins, with special emphasis on Lake Bonneville and the Great Basin, U.S.A.: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 76, p. 189–214, doi: 10.1016/0031-0182(90)90113-L. Currey, D.R., Atwood, G., and Mabey, D.R., 1984, Major levels of Great Salt Lake and Lake Bonneville: Utah Geological Survey Map 73, scale 1: 750,000. Currey, D.R., and Oviatt, C.G., 1985, Durations, average rates, and probable causes of Lake Bonneville expansions, still-stands, and contractions during the last deep-lake cycle, 32,000 to 10,000 years ago, in Kay, P.A., and

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Diaz, H.F., eds., Problems of and prospects for predicting Great Salt Lake levels—Proceedings of a NOAA Conference, March 26–28, 1985: Salt Lake City, University of Utah, Center for Public Affairs and Administration, p. 9–24. Dames & Moore, 1979, Phase II geotechnical and geoseismic study, 220 plus acre parcel, Ensign Downs residential development, Salt Lake City, Utah: Salt Lake City, Utah, Dames & Moore, unpublished consultant’s report, 34 p. Dames & Moore, 1981, Detailed supplemental geotechnical study, Ensign Downs residential development, Salt Lake City, Utah: Salt Lake City, Utah, Dames & Moore, unpublished consultant’s report, 21 p. DePolo, C.M., and Slemmons, D.B., 1998, Age criteria for active faults in the Basin and Range Province, in Lund, W.R., ed., Proceedings volume, Basin and Range Province Seismic-Hazards Summit: Utah Geological Survey Miscellaneous Publication 98-2, p. 74–95. Everitt, B., 1991, Stratigraphy of eastern Farmington Bay: Utah Geological and Mineral Survey, Survey Notes, v. 24, no. 3, p. 27–29. Gilbert, G.K., 1875, Report upon the geology of portions of Nevada, Utah, California, and Arizona, examined in the years 1871 and 1872, in Wheeler, G.M., ed., Report upon geographical and geological explorations and surveys west of the one hundredth meridian, v. III—Geology: Washington, D.C., Government Printing Office, p. 17–187. Gilbert, G.K., 1890, Lake Bonneville: U.S. Geological Survey Monograph 1, 438 p. Giraud, R.E., 2003, Unpublished Utah Geological Survey letter dated August 1, 2003, to Paul Flood, Wasatch-Cache National Forest: Salt Lake City, Utah Geological Survey, 5 p. Gwynn, J.W., 2002, The extraction of mineral resources from Great Salt Lake, Utah—History, development milestones, and factors influencing salt extraction, in Gwynn, J.W., ed., Great Salt Lake—an overview of change: Salt Lake City, Utah Department of Natural Resources Special Publication, p. 201–212. Harty, K.M., and Lowe, M., 2003, Geologic evaluation and hazard potential of liquefaction-induced landslides along the Wasatch Front, Utah: Utah Geological Survey Special Study 104, 40 p., 16 plates. Hernandez, M.W., 2004, A procedural model for developing a GIS-based multiple natural hazard assessment—Case study—southern Davis County, Utah [Ph.D. dissertation]: Salt Lake City, University of Utah, 402 p. Hintze, L.F., 1988, Geologic history of Utah: Brigham Young University Geology Studies Special Publication 7, 202 p. Hylland, M.D., 1999, Comparative evaluation of earthquake sources associated with the liquefaction-induced Farmington Siding landslide complex, northern Utah, in Spangler, L.E., and Allen, C.J., eds., Geology of northern Utah and vicinity: Utah Geological Association Publication 27, p. 203–222. Hylland, M.D., Black, B.D., and Lowe, M., 1997, Geologic hazards of the Wasatch Front, Utah, in Link, P.K., and Kowallis, B.J., eds., Cenozoic to Recent geology of Utah: Brigham Young University Geology Studies, v. 42, Part II, p. 299–324. Hylland, M.D., and Lowe, M., 1998, Characteristics, timing, and hazard potential of liquefaction-induced landsliding in the Farmington Siding landslide complex, Wasatch Front, Utah: Utah Geological Survey Special Study 95, 38 p. International Code Council, 2002, 2003 International Building Code: Falls Church, Virginia, International Code Council, 656 p. Keaton, J.R., 1986, Potential consequences of tectonic deformation along the Wasatch fault: Logan, Utah State University, Department of Civil and Environmental Engineering Final Technical Report prepared for the U.S. Geological Survey: Utah Geological Survey Contract Report 93-8, 23 p., 6 plates. Lin, A., and Wang, P., 1978, Wind tides of Great Salt Lake: Utah Geology, v. 5, no. 1, p. 17–25. Lowe, M., 1993, Hazards from earthquake-induced ground failure in sensitive clays, vibratory settlement, and flooding due to seiches, surface-drainage disruptions, and increased ground-water discharge, Davis County, Utah, in Gori, P.L., ed., Applications of research from the U.S. Geological Survey program, assessment of regional earthquake hazards and risk along the Wasatch Front, Utah: U.S. Geological Survey Professional Paper 1519, p. 163–167. Lund, W.R., 1992, New information on the timing of earthquakes on the Salt Lake City segment of the Wasatch fault zone—Implications for increased earthquake hazard along the central Wasatch Front: Utah Geological Survey, Wasatch Front Forum, v. 8, no. 3, p. 12–13.

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