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 Hi...

Author: Joel L. Pederson | Carol Merritt Dehler

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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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runninghead Contents

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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River

E NG A R

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

Ct Zm Zmf Zb Xlw XWf

WASATCH RANGE

ek

n ee Gr

formation of Dead Horse Pass formation of Red Castle, lower formation of Moosehorn Lake

Zucl Zum

Mount Watson Formation

formation of Red Castle, upper

r

ve Ri

Zud

Zuw

Zucu

formation of Hades Pass

Red Pine Shale

Zur Zuh

TIntic Quartzite

Ct

WESTERN UINTA MOUNTAINS

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Cl

40

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

Zur Zuu Zuj Xr

N

40 O 108O

108O 41 O

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DAY 1 STOPS DAY 2 STOPS

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THRUST AND REVERSE FAULTS (teeth on upthrown side)

NORMAL FAULTS (bar and bar on downthrown side)

UINTA ANTICLINE

Lodore Formation Cl

UTAH

iver te R Whi

iver

Yamp aR

EASTERN UINTA MOUNTAINS

109O

Zuu

SP AR K

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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EXPLANATION

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Roosevelt

Vernal

191

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UTAH

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r

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ve

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Kilometers

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12 16 20

L

8

er

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ork

ry River Strawber

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Duchesne

Zur

er

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Zur

A I ClN S MOUNT

cks Riv

0

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ck

111 O

Zur

6

TA U I N Zuh

109O

COLORADO

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Ct

5

Zud

ork Zud

WYOMING

Flaming Gorge Reservoir Zuj ver Xr 1 Beraeek 2 C 3 4 5 191 Dutch John 3 Sheep Creek BR SR 44 OW Zuu N

Pz

reek

Utah Lake

Heber City

Provo River

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2

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80

H TC A S WA

Salt Lake City

Great 15 Salt XWf Lake

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Uinta Mountain Group

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Neoproterozoic Uinta Mountain Group of northeastern Utah 3

Dehler et al.

4 Miss.

δ13Corg (o/oo PDB) -32 -30 -28 -26 -24 -22 -20 -18 -16

Red Pine shale

Cambrian

fm. of Crouse Canyon?

4

4

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....

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?

3

km

Hades Pass quartzite km

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* ?

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fm. of Crouse 2 Canyon

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mudcracks planar crossbeds trough crossbeds ripplemarks break in section acritarch vase-shaped microfossil

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Moosehorn Lake shale & equivs.

?

hcs

....

base not exposed sh s

c

Uinta Mountain Group Western Uinta Mountains

hummocky crossbeds

reverse-grading ripple cross laminations convolute bedding parallel laminations

km

sh s

Key: Mount Watson Fm & equivalents

Cambrian

organic-rich shale facies sandstone and mixed facies sandstone undifferentiated trough crossbedded facies association low-angle crossbedded facies association sandstone and shale breccia, conglomerate, sandtone, shale metaquartzite *fossil location projected from approximately equivalent and shaley strata westward at Leidy Peak locality

fm. of Outlaw Trail

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

Jesse Ewing Canyon Fm. Red Creek Quartzite

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

c

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

5

M

Deseret Formation

Madison Limestone

360

C 543

Tintic Quartzite

Lodore Sandstone

L NEOPROTEROZOIC

630 ~740

Red Pine Shale

M I D D L E

formation of Hades Pass formation of Crouse Canyon

Mount Watson Formation formation of Dead Horse Pass 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


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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Dehler et al.

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]


9

(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 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 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

Tsl

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 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

A.N. Steely et al.

186

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)


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


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 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.

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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 60

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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 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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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

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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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426

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

P9 Tooele Valley

1.6 N

Tooele Army Depot

P7 1.5

P3

P1

P0

T T T

Bauer

1.3 B5

B8

B6

1.4 B1 South Mountain

Oquirrh Mountains

T B3 Rp Rush Valley

Rg

Stockton

0.5 km

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)

427

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.

H.S. Godsey et al.

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

(0.0)

2.2

(3.5)

A Elevation (m)

P1 1480

P3 P5

P9

1470

100

200

300 400 Distance (m)

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B

P5

1485 Elevation (m)

P7

1475

P7

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400

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800

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1200

Distance (m) C

P1

Elevation (m)

1480

P3

P5 P7

1475 1470

P9

1465 400

800

1200

1600

2000

2400

Distance (m)

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.

B 1550

1450

A

1400

1450

1350

Paleoelevation (m)

1550

1350

1300

1250 24

19

14

9

Age (14 C yr B.P. x 1000)

15.5

15.0

14.5

14.0

13.5

13.0

12.5

Age (14C yr B.P. x 1000)

12.0

11.5

1250 11.0

Paleoelevation (m)

1500

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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430 6.4

(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).


431

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.”

H.S. Godsey et al.

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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.

433

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.

434

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.


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

S

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).

6 83

41°35’

128

9m

g Stockin Molly's

1.6

Directions

5

Little Mountain

1283 m (Appr. historic highstand of Great Salt Lake)

LEGEND

Gilbert shoreline beach deposit or abrasion notch (~1296 m) Smoothed topographic contour Primary spring

5

0

Measured section 1 2 3

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.

436

H.S. Godsey et al.

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).

437

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).

H.S. Godsey et al.

438

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


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

H.S. Godsey et al.

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

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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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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)982.0.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)0252.3.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)0202.3.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.

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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

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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


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).


Figure 4 (continued).

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456

13.1

(21.0)

14.6

(23.5)

14.8

(23.7)

15.5

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18.5 18.9

(29.8) (30.4)

19.0

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21.2

(34.1)

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22.4

(36.0)

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.


458

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


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

STOP 1

Grandstand

Hurricane Fault Figure 7

Relay ramp

#14/2.6/0.03-0.06

STOP 2

Ramp

Basalt flow

#7/4.0/0.12-0.27

Relay

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.

459

STOP 3Moriah

Knoll

#10/2.7/0.18-0.34

STOP 4 #5/4.9/0.33-0.61 #4/7.4/0.05-0.09 #1/4.2/0.28-0.53

1 km

Twin Butte

#2/9.3/0.07-0.13

N

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


460

A

Basalt ridge

Qmb

EF Qyt Qm2

Qm2

Qmb

Qco

F

Qco

Qm2

Pkh

E

Qmb Qco Qm2

Qy1

Qmb

WF

Qm2

Pkh

Ram Pkf p Qy1

Qm2

Qm2

Qyc Qm3

Qm1 Qyc

Qm2

lay

Qm3

Qmb

Pkh

Pt-k Pkf

D

Pkh

Qm2

Pkh

Qyc Pkf

Pkh

Qco

Qco Pkh

C

Base of paleo-canyon

B

A

Pt-k

A'

Qm3

Qm3

A

Qm3

Qy1 Qm3

Qm2

Qm3

Qy1

N

Qm3

C

B Basalt Ridge

Base of paleocanyon

Basalt over relay ramp surface

1600 Elevation, m

(1591 m)

Edge of basalt (1387 m) 5X Vertical Exaggeration A

1

2

1500

1400 3 km A'

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.

41.3

(66.5)

41.7

(67.1)

Road Log

41.9

(67.5)

43.0 45.3 45.5

(69.2) (72.5) (72.9)

Proceed east on the Navajo Trail. Cumulative mi (km) 41.0

(66.0)

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).

Twin Butte

Grandstand

Anderson Junction Shivwits

Escarpment height, m

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

Moriah Knoll Basalt crossed escarpment here

450

461

350

N40E 250

0

N15W

N35W

N25W 20

N15W

N10E

Distance south from Cottonwood Canyon, km

N55E

N20E

40

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).


462

45.8

(73.7)

46.2

(74.4)

47.2

(76.0)

47.5

(76.5)

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

50.3

(81.0)

50.7

(81.6)

51.9

(83.6)

52.9

(85.2)

463

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).


464

A el

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Fault, ball on downthrown side dashed where inferred

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northing (m)

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ann

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1050

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1150

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easting (m)

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315 45

287° ±10° 270

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90

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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.

465

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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Stratigraphic units Young alluvium Post-MRE alluvium/colluvium Post-MRE fissure fill PE (?) colluvium Pre-MRE/PE alluvium

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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).

Elevation, m

18

Elevation, m

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466 L. Amoroso and J. Raucci


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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°]

1

[35°]

[10-12°]

[35°]

11

[47°]

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


468

(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 (

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