OSTRACODES AS INDICATORS OF THE PALEOENVIRONMENT IN THE PLIOCENE GLENNS FERRY FORMATION, GLENNS FERRY LAKE, IDAHO

A thesis submitted to Kent State University in partial fulfillment of the requirements for the degree of Masters of Science

by

Cordelia W. Dennison-Budak

May, 2010

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Thesis written by Cordelia W. Dennison-Budak B.A., College of Wooster, 2007 M.A., Kent State University, 2010

Approved by

______, Advisor Dr. Alison J. Smith

______, Chair, Department of Geology Dr. Daniel Holm

______, Dean, College of Arts and Sciences Dr. Timothy Moerland

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TABLE OF CONTENTS

LIST OF FIGURES………………………………………………………………v

LIST OF TABLES………………………………………………………………vii

ACKNOWLEDGMENTS ...... viii

ABSTRACT ...... ix

INTRODUCTION ...... 1

BACKGROUND ...... 5

Location ...... 5 Regional Geology ...... 10 Western Snake River Plain ...... 12 Geothermal Influences ...... 17

STRATIGRAPHY ...... 19

Idaho Group ...... 19 Glenns Ferry Formation ...... 21 Tuana Gravel ...... 26

PLIOCENE-PLEISTOCENE LAKES...... 27

Black Rock, Utah ...... 27 Bonneville System, Idaho and Utah...... 28 Bear Lake, Utah and Idaho ...... 31 Beaver Basin, Utah ...... 31 Searles Lake, California ...... 32

REGIONAL PALEOECOLOGY OF IDAHO IN THE PLIO-PLEISTOCENE ..35

Pollen ...... 35 Fish ...... 43 Ostracodes ...... 44

METHODS ...... 46

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Field Work ...... 46 Lab Work ...... 47 Isotope Analysis ...... 50 Statistical Analysis ...... 51 Cluster Analysis ...... 52 Principal Component Analysis ...... 53

RESULTS ...... 54

Ostracode Assemblages ...... 54 Limnocythere robusta Assemblage ...... 59 Limnocythere friabilis Assemblage ...... 61 Candona crogmaniana Assemblage ...... 63 Limnocythere ceriotuberosa Assemblage ...... 67 Principal Components Analysis (PCA) ...... 70 Cluster Analysis ...... 77 Ostracode Zonation and Paleolake Phases ...... 77 Stable Isotopes ...... 83

DISCUSSION ...... 88

CONCLUSIONS...... 91

REFERENCES ...... 95

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LIST OF FIGURES

Fig. 1. Map of Hagerman Fossil Beds National Monument ...... 6 Fig. 2. Geologic map of Hagerman, Idaho ...... 9 Fig. 3. Landslide in Hagerman Fossil Beds NM ...... 10 Fig. 4. Map of western and eastern Snake River Plains ...... 12 Fig. 5. Stages of Lake Idaho ...... 14 Fig. 6. Map of western Snake River Plain ...... 15 Fig. 7. Stratigraphy of the Idaho Group ...... 20 Fig. 8. Members of the Glenns Ferry Formation ...... 23 Fig. 9. Hagerman Fossil Beds NM timeline ...... 25 Fig. 10. Map of Pleistocene pluvial lakes in the Great Basin ...... 30 Fig. 11. Map of Searles Lake ...... 33 Fig. 12. Climatic stages of Searles Lake ...... 34 Fig. 13. Pollen chart from the Hagerman Horse Quarry ...... 39 Fig. 14. Mural of Hagerman Fossil Beds NM ...... 40 Fig. 15. Five pollen groups ...... 42 Fig. 16. Plate of Hagerman Fossil Beds NM ostracodes ...... 57 Fabaeformiscandona rawsoni Limnocythere friabilis Ilyocypris cf. bradyi Limnocythere ceriotuberosa Limnocythere robusta Fig. 17. Limnocythere robusta Assemblage ...... 60 Fig. 18. Limnocythere friabilis Assemblage ...... 62 Fig. 19. Candona crogmaniana Assemblage ...... 66 Fig. 20. Limnocythere ceriotuberosa Assemblage ...... 69

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Fig. 21. Principal Component Analysis Variables ...... 74 Fig. 22. Principal Component Analysis Scores ...... 76 Fig. 23. Cluster Analysis ...... 80 Fig. 24. Species vs. Valves/gram of Hagerman Fossil Beds NM ostracodes ...... 82 Fig. 25. Oxygen isotope 18O vs. Elevation ...... 84 Fig. 26. Relative oxygen isotope (18O) and carbon (13C) for Hagerman Fossil Beds NM ostracode species ...... 85 Fig. 27. Oxygen isotope 18O vs. Deuterium ...... 87 Fig. 28. Hypothetical location of Lake Idaho ...... 92

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LIST OF TABLES

Table 1. Previous Works Table ...... 7 Table 2. HAFO Species Table ...... 55 Table 3. Limnocythere robusta Assemblage ...... 59 Table 4. Limnocythere friabilis Assemblage ...... 61 Table 5. Candona crogmaniana Assemblage ...... 64 Table 6. Limnocythere ceriotuberosa Assemblage ...... 68 Table 7. Eigenvalues and Variance of Principal Component Analysis Data ...... 72

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ACKNOWLEDGEMENTS

―What is a dinosaur?‖…‖How can I work with one?‖ These were the first two scientific questions I had ever asked. They were answered by my father when I was 5 years old. He said, ―You could become a paleontologist and study them as a career‖. I said, ―Wow‖.

Ever since that conversation my life has been driven passionately towards one thing, becoming a paleontologist. I have had the support of teachers, friends, colleagues, and family along the way. And I want to thank everyone for their support. But, given the journey‘s long distinctive path I would like to thank publicly those people who have stood out from the rest.

To my parents, Anthony Budak and Jennie Dennison-Budak, and grandparents,

David Dennison and Margaret Dennison, for not simply being there, but actually getting me here. The four of you, through love, compassion, education, patience, humor, and gentle guidance gave me the fundamental tools I needed to achieve my goals and all that life may still have to offer. Faffee and Marga, this is for you.

To my love,Warren Swegal. You have spoken nothing but words of encouragement towards me from day one and have been the most stable support system I hope to ever know. Thank you.

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To my dearest friends, Kristen Davis, Rachel Garzarelli, Carrie Holmes, and

Audrey Speicher, for the long distant cheers of support and love.

To my geology colleagues Kristen Enzweiler and Jenna Hojanowski for their willingness to share stories and understand my geological problems.

And to my fearless advisor, Dr. Alison Smith who led me through this with her head held low, holding my hand every possible step of the way. THANK YOU!

To my ‗field‘ advisor, Dr. Donald Palmer, for teaching me the best way to dig for ostracodes and follow the stratigraphy whether or not I thought it was right.

Also, none of this would have been possible without help from the Katherine

Moulton Scholarship and Graduate Student Senate (GSS) Grant for their financial support. For the cooperation of Hagerman Fossil Beds National Monument including

Phil Gensler, Bob Lorkowski, and my wonderful field crew, Amberly Lynott, and Sarah

Heinemann, I thank you: you were all invaluable to this project. And a special thanks to the National Park Service for hiring me to do the work I enjoy the most.

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ABSTRACT

Lake Idaho was a large and permanent lake of Pliocene age located in the southwest portion of what is now Idaho and parts of eastern Oregon. This project looks at ostracodes as paleoenvironmental indicators from an extension of Lake Idaho, the Plio-

Pleistocene Glenns Ferry Lake, located in Hagerman Fossil Beds National Monument

(HAFO), Hagerman, Idaho. Fluvial-lacustrine sediments from the Glenns Ferry

Formation dominate HAFO‘s geologic record. Deposition of Lake Idaho began approximately 3.8 Ma with the deep-lake phase beginning 3.5-3.3 Ma and lasted until around 2.4 Ma (early Pleistocene time). Previous studies indicate that during the

Pliocene, ENSO was in a permanent El Niño phase. However, based on cluster analysis, principal component analysis, and isotopic 18O analysis of the ostracodes collected from

HAFO, I demonstrate that ENSO was already cycling between La Niña and El Niño phases when HAFO lacustrine sediments were deposited in the late Pliocene through early Pleistocene. The range of 18O and corresponding 2H values is consistent with the range of modern winter and summer precipitation values, indicating a precipitation pattern similar to that of today. This conclusion is also supported with multi-proxy evidence from pollen and fish fossils, collected from the literature. Four ostracode assemblages suggest four paleolake phases. The first and earliest in this study, the

Limnocythere robusta Assemblage represents a fresh but variable environment. The

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second, the Limnocythere friabilis Assemblage represents an oligohaline environment.

The third, the Candona crogmaniana Assemblage represents a second fresh but variable environment. The fourth and final assemblage, dominated by Limnocythere ceriotuberosa represents a final oligohaline environment. The ostracode assemblages, the paleolake phases and the initiation of a cyclic ENSO phase show the start of the modern climatic variability in Idaho in late Pliocene time.

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Introduction

Hagerman Fossil Beds National Monument (HAFO) in south central Idaho is primarily well known for its superb vertebrate fossil collections representative of the Late

Pliocene and Early Pleistocene of western North America. The fossils present in the area are there because of Plio-Pleistocene Lake Idaho. HAFO‘S geologic record is dominated by Lake Idaho‘s sediments, which are part of the Glenns Ferry Formation. However, considering HAFO‘s part in the Plio-Pleistocene sedimentary fluvio-lacustrine record of south central Idaho is only part of understanding its role in a larger paleoenvironmental reconstruction. Determining the evolution of the El Niño Southern Oscillation (ENSO) cycle is key to understanding the paleoenvironments of western North America of Plio-

Pleistocene age and why the fossils and sediments of HAFO are found as part of a large

Plio-Pleistocene lake complex known as Lake Idaho.

Deposition of the Glenns Ferry Formation began approximately 3.8 Ma, with the deep-lake phase beginning 3.5-3.3 Ma, and lasted until around 2.4 Ma (Smith, 1987;

Thompson, 1996). The depositional environment at HAFO is of shallow, highly sinuous meandering streams with sandy deposits and by muddy flood plain deposits from standing water (Lee et al., 1995; McDonald et al., 1996).

Lake Idaho was present during the majority of the Pliocene because of higher than modern Pacific sea surface temperatures (SST) that formed high amounts of atmospheric

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precipitation from the evaporated water (Bartoli et al., 2005). Also, because the Rocky

Mountain summit elevations were 1000 meters lower, the precipitation was delivered further inland (Bonham et al., 2009). El Niño was considered to be in a ‗permanent‘ phase at this time and is thought to be the primary reason for these higher than modern

Pacific SST and increased precipitation that reached western North America during the

Pliocene (Bonham et al., 2009; Kukla et al., 2002).

El Niño is defined as an event when ―lower air pressure in the east weakens the westward atmospheric pressure gradient leading to unusually high SST extending across the central and eastern tropical Pacific Ocean and enhanced convection over the area‖

(Bonham et al., 2009, p. 128). During the Pliocene, El Niño is thought to be a primary feature of the climate that helped keep the western coast of North America steadily warm

(Bonham et al., 2009; Huybers and Molnar, 2007; Kukla et al., 2002). By 2.6 Ma, the

Isthmus of Panama had closed and climate began to change (Lunt et al., 2008). The closing led to decreased water mass mixing, the initial formation of the North Atlantic

Deep Water (NADW) and the strengthening of the North Atlantic thermohaline circulation, increased sea surface temperatures along the equatorial Pacific which led to increased evaporation in the North Atlantic, higher precipitation on the Northern

Hemisphere at higher latitudes and the final intensification of the Northern Hemisphere

Glaciation (Bartoli et al., 2005; Bonham et al., 2009; Huybers and Molnar, 2007; Kukla et al., 2002; Lunt et al., 2008). With all these changes, ENSO started to cycle between an

El Niño and La Niña phase (Lunt et al., 2008). One of the outcomes of this study is to present evidence of the timing of the change from a permanent to cyclic ENSO.

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The purpose of this project was to determine paleoenvironmental conditions of

Hagerman Fossil Beds National Monument though the use of ostracodes as indicators.

The primary goals of this project were to determine the following:

1. Is the lacustrine record at HAFO consistent with other lacustrine records of

similar age in the region?

2. Can ostracodes help constrain the temperatures, hydrochemistry, and age of

the deposits at HAFO?

3. What are the 18O isotopic ratios in the ostracode carapaces in the HAFO

sediments?

4. Does HAFO record a lacustrine history of a change in seasonality or El Niño

conditions?

During the summers of 2007 and 2008, the collection of 75 samples yielded 15 common ostracode taxa, which provided enough information to successfully answer these questions and ultimately determine four lake stages of HAFO during the early

Pleistocene, that were from a fresh but variable lake to an oligohaline environment. We also determined that the ostracode paleoecological and isotopic record indicate HAFO sediments were deposited in the final phases of Lake Idaho, and after ENSO started to vary between an El Niño and La Niña phase.

Since the beginning of this project in 2007, a new timescale has been applied to the Pliocene-Pleistocene boundary. The previous timescale indicated that based on

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cooling trends, the Pliocene-Pleistocene transition began around 1.8 Ma (Gibbard and

Cohen, 2009). It has now been established, that based on those same cooling trends, the new boundary is set at 2.6 Ma (Gibbard and Cohen, 2009).

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Background

Location

The location for field research was conducted at the Hagerman Fossil Beds

National Monument (HAFO) in Hagerman, Idaho. The town of Hagerman is located in south-central Idaho on the eastern shore of the Snake River as it cuts north, while HAFO makes up 4,400 acres along 5 miles of the western shore (Figs. 1 and 2). The Glenns

Ferry Formation and some of the Tuana Gravel of the Idaho Group located within the western Snake River Plain (WSRP) make up HAFO‘s depositional units (Thompson,

1996).

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Fig. 1. Map of highlighted (green) range of Hagerman Fossil Beds National Monument and Hagerman, Idaho (Sun Valley Guide, 2003).

In 1988, the Hagerman Fossil Beds became a National Monument and the

Hagerman Horse, Equus simplicidens, became Idaho‘s state fossil. Most famous for the

Hagerman Horse, no other park has preserved such variety, quantity, and quality of fossils from the Pliocene than HAFO (Multiple, 2002). Today, HAFO has produced

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more than 180 invertebrate and vertebrate species and 35 plant species from the fossil record. The Monument is also one of three National Parks to contain sections of the

Oregon Trail National Historic Trail, which passes through the park property.

I chose to study ostracodes because they can be very useful in tracking changes in the positions of ocean water masses, tidal ranges, salinity, solute source and history, and water depth (Smith and Horne, 2002). However, it is also important to consider other types of proxies while determining paleoenvironments as one proxy may disregard qualities another recorded. The following is a list of publications that can be found as previous works of reconstructed paleoclimatic and paleoecological aspects of HAFO

(Table 1).

Table 1. Table of other proxies and suggested readings used in the paleoclimate and paleoecological reconstruction of HAFO during the Plio-Pleistocene.

Previous Works Table Proxy Citation Fish Smith, 1975; Smith, 1987; Smith and Patterson, 1994. Microtine Repenning, 1985; Ruez and Gensler. 2008; Zakrzewski, 1969 Rodents Ostracodes Forester, 1985; Forester, 1991; Swain 1986 a and b; Thompson, 1996 Pollen Leopold and Wright, 1985; Thompson, 1996 Vertebrate Bjork, 1970; Brodkorb, 1958; Fine, 1964; Ford and Murray, 1967; Fossils Gazin, 1936; Samuels et al., 2009; Shotwell, 1961

Recently, geologic activity in HAFO has included major landslides. The geology of the area is mostly composed of lacustrine deposits like sand, which can have a high hydraulic conductivity (Forester, 1991; Lee et al., 1995; Malde, 1991; McDonald et al.,

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1996; Sadler and Link, 1996; Thompson, 1996). As groundwater delivered by the Bell

Rapids Irrigation Company, to the top of the bluff percolates through the sediment it often reaches a sand bed incased in mud and forms a perched aquifer (Lee et al., 1995).

These perched aquifers leak on the slope face in HAFO forming vegetated areas not previously covered. As the sediment becomes saturated, the unstable slope often gives way to a landslide (Fig. 3). These landslides are detrimental to the park for a variety of reasons including (but not limited to) the displacement of fossil localities and the reburial of exposed ones as well as the hazards, associated with landslides, to people and animals.

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Fig. 2. Clipping from the geologic map of the Hagerman quadrangle, Gooding and Twin Falls counties, Idaho (Othberg et al., 2005).

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Fig. 3. Photograph of recent landslide showing Pliocene lakebeds on south end of monument along the Snake River. (Photo courtesy Alison J. Smith, 2007).

Regional Geology

The Snake River Plain is composed of two main halves, the western and the eastern. The western Snake River Plain (WSRP) is a structural basin filled with sedimentary deposits above a basalt slab (Malde, 1991). The eastern Snake River Plain

(ESRP) is composed primarily of rhyolitic volcanic rocks and basaltic lava flows (Malde,

1991). The basalt slab in the WSRP that Malde (1991) refers too was most likely emplaced during basin rifting perpendicular to the eastern Snake River Plain trend. The

ESRP is composed mostly of buried rhyolitic calderas covered with basalt that formed

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during the eastern migration of the Yellowstone-Snake River Plain hotspot over the past

17 Ma (Lee et al., 1995; Malde, 1991; McDonald et al., 1996). To the southwest of the

ESRP, rhyolite and basalt form an upland called the Owyhee-Humboldt Plateau (Fig. 4)

(Malde, 1991; McDonald et al., 1996).

The SRP forms a crescent shape about 70 km wide in its western part, 90-100 km in its eastern and a total of approximately 600 km long (Fig. 4) (Malde, 1991). The western SRP (WSRP) and eastern Snake River Plain (ESRP) are separated by their stratigraphy and their distinctive ages (Malde, 1991). The WSRP includes HAFO and, thus, where I will focus the discussion of my geologic descriptions.

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Fig. 4. Map of the western and eastern Snake River Plains, both in a bolded outline. The grey stippled area represents the eastern Snake River Plain described in the text. The red star represents approximate location of Hagerman, ID. (Modified from Malde, 1991, p. 252).

Western Snake River Plain

The western Snake River Plain (WSRP) is described as a fault-bounded basin with uplands to the north and south, and filled with mostly sedimentary and minor volcanic units, that grew towards the southeast until the Late Pliocene time (~16 to ~3

Ma) (Lee et al., 1995; Malde, 1991). The WSRP is approximately 260 km long and narrows eastwardly from 70 km, at its widest, to 35 km wide (Malde, 1991; Sadler and

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Link, 1996). Its altitude varies along its length, but in HAFO it reaches up to 3,400 m above sea level (Lee et al., 1995; Malde, 1991). Near Hagerman 1.7 km of lacustrine sediments are underlain by basalt and, further down, tuffaceous sediments and silicic and intermediate volcanic rock (Malde, 1991).

Sadler and Link (1996) have grouped the WSRP into four depositional or geologic phases: Initial Basin and Range Phase (17.5-14 Ma), Hot Spot Shoulder Phase

(~14-9 Ma), Lake Idaho Phase (~9-2 Ma), and Longitudinal Drainage Phase (~2 Ma-

Recent) (Fig. 5) (McDonald et al., 1996). The Initial Basin and Range Phase began when rifting created the Ore-Ida graben along Basin and Range normal faults (Fig. 5) (Sadler and Link, 1996). The volcanic rock that lies at the base of WSRP is thought to come from this stage of initial rifting, 17 Ma (Kimmel, 1982; Malde, 1991). The WSRP is also low in elevation compared to the surrounding areas, reflecting this isostatic compensation for the basalt (Malde, 1991).

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Fig. 5. The four stages of the development of the western Snake River Plain, as interpreted by Sadler and Link, 1996.

Between 14-9 Ma, both Malde (1991) and Sadler and Link (1996) note local rhyolitic eruptions and regional thermal uplift in the WSRP. Both agree these rhyolitic volcanic rocks came from multiple volcanic centers including the Owyhee-Humboldt and

Bruneau-Jarbridge volcanic centers (Malde, 1991; Sadler and Link, 1996). Sadler and

Link (1996) refer to this event as the Hot Spot Shoulder Phase where regional uplift and subsequent subsidence occurred on the southwestern portion of the WSRP graben from thermal inflation above the Yellowstone hot spot (Figs. 4 and 5). Sadler and Link (1996)

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note that the WSRP is located on the northern flank of the thermal bulge, not in the middle, where sediment accumulation is scarce and therefore not apparent in the depositional record (Fig. 6) (Sadler and Link, 1996).

Fig. 6. Map of the western Snake River Plain with Hagerman near the bottom right of the Plain. The red star represents Hagerman, ID. (Modified from Malde, 1991, p. 257).

The route by which the hot spot traveled (or more correctly, the route by which the lithospheric plate moved over the hotspot) is referred to as the eastern Snake River trend (Fig. 4) (Malde, 1991). From around 9-2 Ma the hot spot moved in an easterly direction, lending to ―detumescence of the thermal hot spot bulge‖ causing further

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subsidence of the WSRP into a deep basin (Lee et al., 1995; Sadler and Link, 1996, p.

46). It is during this time, which Sadler and Link (1996) refer to as the Lake Idaho

Phase, when sediments found in HAFO (i.e. the Idaho Group and Glenns Ferry

Formation) were deposited. Kimmel (1982) suggests that volcanism in the west may have created the lakes by blocking outlets in the early Pliocene. The volcanism that

Kimmel refers to is considered to be the rhyolitic Idavada Volcanics, which are overlain by the Banbury Basalt (Kimmel, 1982; Malde, 1991). Thompson (1996) discusses an estimated time-line for Lake Idaho starting with the Glenns Ferry Formation deposition prior to ~3.8 Ma through ~2 Ma. He suggests the deep-lake phase of Lake Idaho began around 3.5 or 3.3 Ma until 2.5-2.4 Ma (Thompson, 1996).

Considered part of the eastern Snake River Plain trend, the Idavada Volcanics

(dated around 10-11 Ma) and Banbury Basalt, in the WSRP, make up the Mount Bennett

Hills in the north and the Owyhee Plateau in the south (Fig. 6) (Malde, 1991). There is an additional basalt that is not connected to either the Idavada or Banbury deposits, east of Salmon Falls Creek and outside the WSRP boundary (Fig. 6), which appears below

Quaternary aged basalts of the ESRP and above the Banbury Basalt (Malde, 1991). This basalt is interbedded with the Glenns Ferry Formation, and has been estimated to be approximately 4 to 5 Ma in age (Malde, 1991). However, zircons in an ash bed (Peters

Gulch Ash) below the interbedded basalt are dated 3.75  0.36 Ma (Malde, 1991).

The Longitudinal Drainage Phase represents the final spillover (or capture by the

Columbia-Salmon River drainage) of Lake Idaho to the north at Hell‘s Canyon (Malde,

1991; Sadler and Link, 1996). The 2 Ma Tuana Gravel deposit indicates the beginning

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stages of the drainage of the Snake River. Deposition of the Tuana Gravels occurred when the Snake River existed as a high-energy environment and drainage was, presumably, unobstructed (Malde, 1991).

The Longitudinal Drainage Phase, is depicted by Sadler and Link (1996), as successive stages of high-energy fluvial sedimentation graduating into lower energy regimes with short-lived drainage obstructions. Regression and spillover of Lake Idaho‘s outlet lead to an upward-fining, pulsed, deposition of the Tuana Gravel followed by the lacustrine Bruneau Formation (Sadler and Link, 1996). Malde (1991) suggests that four successive canyons, each blocked by canyon-filling basalts, led to four different river channels. The last of which is represented by the Snake River along its current path. The most recent lava dam (stage 4) diverted the ancestral Snake River to the northeast and into its modern path, because overflow along the southwestern side was blocked by the

Owyhee Mountains (Malde, 1991). Sadler and Link (1996) make reference to climatic changes of the Pleistocene, such as increased stream discharge, as causation for spillover.

This is discussed later, in the sections on paleoecology and paleoclimate.

Geothermal Influences

The WSRP also contain part of southern Idaho‘s hydrothermal resources (Malde,

1991). Southern Idaho takes full advantage of these thermal resources, as more than 250 houses in the Boise Warm Springs Water District are heated with geothermal water, from aquifers that have been heated by magma. The temperature of the springs averages 77C,

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but ranges from 40C to 83C (Malde, 1991). In contrast, shallow aquifers in the sedimentary deposits rarely pass 35C. Deuterium and oxygen-18 in the water suggest that recharge occurred when the climate was 3 to 5C colder than present, with a residence time of 3,400 to 6,800 years (Malde, 1991).

I found hot and cold springs located sporadically throughout the Hagerman

Valley, like Thousand Springs (cold), Banbury Hot Springs, and elsewhere. For the public as well as private use, these springs are also utilized as alligator farms. The alligators are primarily farmed for their meat and hides and maybe a pet or two.

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Stratigraphy

Idaho Group

Stratigraphy in the Hagerman Fossil Beds NM are primarily sedimentary (and mostly lacustrine), but some basalt is interbedded in several parts of the WSRP (Malde,

1991; Thompson, 1996). For example, both Kimmel (1982) and Malde (1991) find sedimentary deposits of the Lake Idaho Phase interbedded with basalts in the Glenns

Ferry Formation. Progressing eastward, the sedimentary deposits decrease in age from about ~16 Ma to 1 Ma in age (Malde, 1991). These deposits are lumped together into the

Idaho Group, which ranges from ~11-2 Ma (Fig. 7) (Malde, 1991).

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2.6 Ma

Fig. 7. Stratigraphy of the late Cenozoic Idaho Group in the western Snake River Plain. The blue-shaded area represents HAFO‘s primary sedimentary deposits including part of the Tuana Gravel. The red arrow represents the new geologic time change between the Pliocene and Pleistocene. (Gibbard and Cohen, 2009; Modified from Malde, 1991, p. 258).

The Idaho Group was deposited in lakes, floodplains, and streams that were affected by local basaltic volcanism and subsidence of the WSRP graben as well as silicic volcanism from the ESRP (Lee et al., 1995; Thompson, 1996). The Idaho Group of the

WSRP, is defined as a collection of seven formations from the Miocene to Quaternary that includes the Poison Creek, Chalk Hills, Glenns Ferry, Tuana Gravel and Bruneau

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Formations (Fig. 7) (Kimmel, 1982; Lee et al., 1995; Malde, 1991; Sadler and Link,

1996; Thompson, 1996). The Banbury Basalt near the base of the group primarily refers to the eastern Snake River Plain trend (Lee et al., 1995; Malde, 1991).

Glenns Ferry Formation

The Glenns Ferry Formation is an assemblage of floodplain, lake and stream deposits interbedded with lava flows that occupy an area of several thousand square miles in the WSRP (Kimmel, 1982; Leopold and Wright, 1985; Malde, 1972). Deposits of

Glenns Ferry Formation are up to 2000 ft thick (Thompson, 1996). Deposition began approximately 3.8 Ma with the deep-lake phase beginning 3.5-3.3 Ma and lasted until around 2.4 Ma (Smith, 1987; Thompson, 1996).

The Glenns Ferry Formation makes up most of the Hagerman Fossil Beds NM.

Near the center of the WSRP, the GFF contains sediments estimated around 2 Ma, the early Pleistocene (Forester, 1991; Kimmel, 1982; Malde, 1991; Ruez, 2006; Thompson,

1996). As the Glenns Ferry Formation thins towards the eastern border of the WSRP, it reaches its youngest ages in this area (Malde, 1991). It finally pinches out approximately

20 km east of Hagerman between the Banbury Basalt and overlying Tuana Gravel (Lee et al., 1995; Malde, 1991). About 183 m (604 ft) of the Glenns Ferry Formation is exposed at HAFO with elevations ranging from 853 m to 1,037 m (~2,800 to ~3,400 ft) (Lee et al., 1995). The type section of the Glenns Ferry Formation contains lacustrine and fluvial sediments with interbedded basaltic lava flows (McDonald et al., 1996).

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The depositional environment of the Glenns Ferry Formation at HAFO is predominantly of shallow, highly sinuous meandering streams with sandy deposits encased by muddy flood plain deposits from standing water (Lee et al., 1995; McDonald et al., 1996). Paleocurrent measurements suggest these streams flowed to the north and northwest (McDonald et al., 1996). Tracings of sand bodies, suggests the southern end of the monument was nearer the main channel of the ancestral Snake River (McDonald et al., 1996).

McDonald et al. (1996) note that the sedimentary deposits were flood plain or lacustrine deposits with little lateral migration. Smith (1987) describes the deep lake sediments of Lake Idaho as massive, pale silts lacking laterally extensive laminar structures. Based on the vertical stacking of these deposits, it is suggested that the rate of sedimentation was the same as the rate of the basin subsidence (Lee et al., 1995;

McDonald et al., 1996). Based on the lack of chemical precipitates in the lacustrine sediments, Thompson (1996) suggests Lake Idaho must have had an outlet. He further states that fossil fish and snails indicate that the outlet drained west, eventually into the

Pacific Ocean (Thompson, 1996).

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Fig. 8. Three generalized stratigraphic members of the Glenns Ferry Formation in HAFO. ASL – Above sea level. Notice the two ash layers, the Peters Gulch Ash and the Fossil Gulch Ash, which are discussed later. (Modified from McDonald et al., 1996, p. 23).

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McDonald et al. (1996) and Lee et al. (1995) also point out that, at Hagerman, the

Glenns Ferry Formation has three members (Fig. 8). The lowest member, a lower flood plain environment, is about 67 meters thick and capped by the Peters Gulch (rhyolitic) ash (McDonald et al., 1996). The middle member, an interval of marshy flood plain deposition, is about 30 meters thick and lies between the Peters Gulch ash and higher

Fossil Gulch (dacitic) ash (McDonald, 1996). The last member, an upper flood plain, is

91 meters thick and continues from Fossil Gulch ash up towards the unconformable contact with the Tuana Gravel, this member also includes the Horse Quarry (McDonald et al., 1996). It is within all these members, that this research focuses on describing four environmental stages based on ostracode assemblages gathered from HAFO (Fig. 9).

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Fig. 9. Figure showing the most recently established ages of two identified beds within HAFO. The Horse Quarry date of 3.19 Ma is an interpolated date. The Fossil Gulch Ash is dated at 3.8 Ma and the Peter‘s Gulch Ash at 3.93 Ma. The last appearance of L. robusta at ~2828‘ represents the final stages of a deep-water phase (Composed from field notes and Hart and Brueseke (1999)).

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Tuana Gravel

The Tuana Gravel, crops out above the Glenns Ferry Formation at approximately

3300 feet in HAFO. It is considered stratigraphically stable, with no faulting (Malde,

1991). It represents at least 10 to 60 m of silt and mud, sand, and gravel (Sadler and

Link, 1996). About 20 m of the Tuana Gravel is actually within HAFO, best seen on the northern portion of the Monument (Sadler and Link, 1996). The Tuana Gravel, as mentioned above, marks the final Snake River canyon stage—marking the ends of the

Glenns Ferry Formation and the Lake Idaho Phase (Malde, 1991; Sadler and Link, 1996).

Sadler and Link (1996) identified two primary lithofacies associations: gravel/sand and fine sand to mud. The gravel and sand lithofacies represents a high-flow deposit, next to permanent, low-flow channels in a braided stream (Sadler and Link, 1996). While, in contrast, the fine sand and mud lithofacies represents flood plains with a low flow regime

(muds) and waning overbank flows (sands) (Sadler and Link, 1996). The majority of finer deposits represent laterally migrating channels and occasional flooding (Sadler and

Link, 1996). Based on paleocurrent data, Sadler and Link (1996) interpret the Tuana

Gravel as being deposited by a laterally migrating braided stream.

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Pliocene-Pleistocene Lakes

During the Pliocene and into the early Pleistocene, the western United States was capable of supporting large, permanent, freshwater lakes like those of the Great Lakes today in the east. This section describes some examples of those lakes, other than Lake

Idaho, in which ostracode research was conducted that showed some similarities to Lake

Idaho‘s paleoecology.

Black Rock, Utah

Ostracodes found in a core collected from Black Rock, Utah were used to help describe the presence of a now-extinct lake. Based on the ostracode assemblages, the lake is suggested to have experienced at least three main phases between Mid- Pliocene to Mid-Pleistocene time (Thompson et al., 2009). The first phase involved a calcium- enriched, shallow, saline, wetland, based on the presence of Limnocythere staplini

(Thompson et al., 2009). The second phase replaces L. staplini with Elkocythereis bramlettei, which implies a shift in source water indicated by a carbonate-enriched environment and thus an increased stream input (Thompson et al., 2009). The researchers involved are unclear at this time whether the increase in stream input is from an increase in precipitation. The third phase replaces E. bramlettei with L. ceriotuberosa, suggesting the water chemistry did not change, but that the lake got deeper and thus

27 28

colder or salinity levels were above E. bramlettei’s tolerance level (Thompson et al.,

2009). L. ceriotuberosa was also found at HAFO, presumably from the extinct Lake

Idaho. Forester (1991, p. 141) suggested this presence represented a ―fresh to saline eurytopic lake‖.

Bonneville System, Idaho and Utah

Another example of ostracode assemblages includes those collected from five deep cores recovered from the Bonneville System (now Great Salt Lake) (Fig. 10). Lake

Bonneville had a surface area of around 51,300 km2, a volume of 9,500 km3, and was over 370 m deep (Lowe and Walker, 1997). From prior paleolimnologic work conducted on the Great Salt Lake Basin, Kowalewska and Cohen (1998) confirmed two major basins called the North Basin and South Basin. The two most common ostracode species were F. rawsoni and Cyprideis beaconensis, which can be seen, respectively, today in marshes and as a lacustrine species (Kowalewska and Cohen, 1998) in western North

America. Ostracode data showed that laterally and/or vertically though time, the North

Basin existed as a fluvial environment, lake, saline lake, playa, and marsh since 5.0 Ma

(Kowalewske and Cohen, 1998). The South Basin, however, was very different, and showed a dominant fluvial environment with a later development of ponds, playas, marshes, and lakes (Kowalewske and Cohen, 1998). Expansion of the lake system could have been controlled by other systems like Lake Idaho, which at the time is thought to have been at a higher elevation relative to the Great Salt Lake Basin (Kowalewske and

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Cohen, 1998). Based on the paleolimnologic evidence, Kowalewske and Cohen (1998) suggest a precipitation increase occurred during the early Pliocene. Forester (1991) agrees that a high precipitation is needed to sustain large lakes, like Lake Idaho, and that at 3.5-2.5 Ma the climate was wetter than it is today.

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Fig. 10. Pluvial lakes in the Great Basin at their maximum Late Pleistocene extent. Notice the difference in size between the Pleistocene Bonneville Lake and the modern Great Salt Lake. Red star indicates approximate location of Hagerman, ID (Modified from Reheis and Bright, 2008, p. 2).

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Bear Lake, Utah and Idaho

Bright et al. (2006) was able to interpret the hydrologic budget for Bear Lake

(Idaho-E Utah) over the last 250,000 years. Bear Lake solutes are dominated by calcium, magnesium, and bicarbonate (Bright et al., 2006). When the lake is topographically closed, most of the calcium is lost to carbonate production (Bright et al., 2006).

Ostracodes, which were studied for their 18O and 13C isotopes were collected from the core and included Cytherissa lacustris and Candona sp. Based on the isotopes of the ostracodes and bulk sediment samples, Bright et al. (2006) concluded that the bulk

18 sediment had less of a response to the changes in  Olake than the ostracodes, which responded greatly to isotopic ranges throughout the lake‘s history. The ostracodes

18 showed a  Olake range from -17 to -5‰ most likely consistent with large glacial- interglacial fluctuations, found in other proxy climate records, over the past 250,000 yr.

(Bright et al., 2006).

Beaver Basin, Utah

Beaver Basin is located in south-central Utah and contains Pliocene and

Pleistocene lacustrine sediments similar to those found in HAFO (Forester and Bradbury,

1981). The study looked at uranium that had leached from nearby volcanic source rocks and the influence it had on water chemistry. Forester and Bradbury (1981) determined water chemistry based on ostracode and diatoms and established four lacustrine systems or depositional environments within Beaver Basin, UT. The water chemistry data was

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based on nine localities, which included at least thirty-two ostracode-bearing assemblages in the area (Forester and Bradbury, 1981). Forester and Bradbury (1981) also came across L. robusta, just as I had in HAFO, and noted that L. robusta did not occur again after the Huckleberry Ridge Ash, which is dated to be 2 Ma (Forester and Bradbury,

1981).

Searles Lake, California

Not all paleoenvironmental reconstructions have to come from biological proxies, some can just be from sediment cores. Searles Lake, now a closed basin, in the Great

Basin, just southwest of Death Valley, California, began at 3.2 Ma in the Searles Valley, and was the fourth basin downstream of six basins (Fig. 11) (Smith, 1984). Based on the sediment record of a 930 m core collected from the center of the lake, Smith (1984) showed that Searles Lake received its water from a succession of overflowing lakes upstream, originally coming from tributary runoff from the Sierra Nevada. This overflow cascade began when the headwaters of the Owens River, specifically the San Joaquin

River was blocked by a volcanic flow and the overflow was rerouted east of the Sierra

Nevada in the Searles Valley at 3.2 Ma (Smith, 1984).

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Fig. 11. Map showing location of Searles Lake (shaded blue), CA (Modified from Smith, 1984, p. 2).

From the Pliocene through to the Holocene, Searles Lake experienced two wet, three intermediate, and two dry regimes (Fig. 12) (Smith, 1984). However, none of the three types of regimes appears to indicate any relation to global changes. Instead, the global glacial and interglacial stages indicated by marine isotope stages records were only apparent in the intermediate phases, and not the wet or dry regimes of the lake (Smith,

1984). Based on this information, Searles Lake is considered one of the better indicators of regional aridity or high moisture content, but not global climate (Smith, 1984). As an indicator of regional climate, Searles Lake‘s sedimentation rates matched those of water runoff or moisture content. Whenever there were low sedimentation rates, there appeared

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to be low moisture (Smith, 1984). These rates, compared to the other previously mentioned lakes, share regional climatic patterns seen in Lake Idaho.

These lakes record the ‗outcome‘ of a different climate state in western North

America than observed today.

Fig. 12. Figure showing the two wet, three dry, and four intermediate stages of Searles Lake, CA during the newly established Pliocene-Pleistocene transition time (Smith, 1984, p. 3).

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Regional Paleoecology of Idaho in the Plio-Pleistocene

The Middle Pliocene is considered one of the last sustained global warm periods prior to the onset of the Northern Hemisphere Glaciation at the beginning of the

Pleistocene (Thompson, 1996; Thompson and Fleming, 1996). One of the major differences between the distribution of modern and Pliocene vegetation is its far reach into ―higher latitudes of both hemispheres‖ (Thompson and Fleming, 1996, p. 44).

Coniferous forests used to grow in current areas of ―polar desert‖ or tundra, making the change from the Pliocene into the Pleistocene a clear overall shift into a colder modern climate (Thompson and Fleming, 1996, p 44).

Paleoecology looks at the way in which past organisms interact with other members of different species and with the ―physical and chemical environment‖ (Cronin,

1999, p. 94). Biological proxies are used to help in the reconstruction of these interactions. The use of proxies like pollen, fish, and ostracodes are discussed here to explain and describe the paleoenvironment of Hagerman, Idaho during the Pliocene.

Pollen

Pollen is used as a proxy in Hagerman, as well as other locations, because of its association with vegetation and paleoclimate records within a region (Cronin, 1999).

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Depending on the dominance of one species over to another, an overall estimate can be made about the local area‘s vegetation at the time of deposition.

Determining the modern climate pollen spectra is important, as well as comparing it to the fossil record in a particular area. Leopold and Wright (1985) conducted a study of pollen at the Hagerman Horse Quarry located within the Monument. Based on modern pollen samples and instrumental data of Hagerman and surrounding areas, winter precipitation is 2-3 times greater in forested areas than steppe (Leopold and Wright,

1985). Also, evaporation is twice the rate of precipitation in the steppe areas, like

Hagerman (Leopold and Wright, 1985). In the fossil pollen record of the Glenns Ferry

Formation at HAFO (below the Hagerman Horse Quarry), there are at least three distinct zones of vegetation (Fig. 13) (Leopold and Wright, 1985).

In the lower part of the Glenns Ferry Formation pollen record at HAFO, there is a dominance of Pinus with lesser amounts of Artemisia (Leopold and Wright, 1985). This type of sample corresponds to a modern mixed conifer forest above 8000 feet (Leopold and Wright, 1985). The middle part of the Glenns Ferry Formation pollen record at

HAFO is represented by lower values of Pinus and higher levels of grasses along with

Artemisia and relative higher level of Juniperus (Leopold and Wright, 1985). The upper part shows ―sporadic peaks of grass pollen‖, high Pinus levels, Artemisia, and small amounts of Juniperus (Leopold and Wright, 1985, p. 337). Also found near the upper portion were aquatic-based plants thought to be from the floodplain (Leopold and Wright,

1985). Leopold and Wright (1985) summarized the area as having a low annual precipitation with cooler temperatures as well as a pine woodland or open forest

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vegetation with diverse steppe elements. Leopold and Wright (1985) surmise that the

Smithsonian Institute was wrong in their description of the Hagerman mural (Fig. 14) and suggest rather the Snake River was surrounded by a ―woodland or open forest where…grassy understory filled the openings‖ (p. 340).

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Fig. 13. Chart of the pollen recovered from the Hagerman Horse Quarry in the Glenns Ferry Formation. Distance is in feet from first ash layer at left and not in elevation from sea level or the Snake River. I have highlighted the two pollen types that are addressed within this paper, Ambrosia and Pine. Sagebrush appears black and undetermined in white (Modified from Leopold and Wright, 1985, p. 336).

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40

Fig. 14. Copy of Jay Matternes Mural of HAFO during the Pliocene (NPS.gov).

In a study performed by Thompson (1996), the use of pollen was applied to determine environmental conditions surrounding the Glenns Ferry Formation or Lake

Idaho location during the middle Pliocene. In Thompson‘s (1996) research, the pollen recovered from 79 Glenns Ferry samples indicated five zones (of depth) each originating from a large source area (probably from the WSRP) (Thompson, 1996). Within the five zones of the Glenns Ferry deposit (G1-G5), each pollen type preferred particular temperatures. For example, Ambrosia-type (ragweed and relatives) prefers warmer climates, while Pinus (pine) and Artemisia (sagebrush) prefer cooler ones (Thompson,

1996). Artemisia, in particular, is negatively correlated with magnetic susceptibility, a feature not totally understood by Thompson (1996). Although, low Artemisia values

41

(with high MS values) may indicate a climatic state during transitions from moist cool conifer climates to warm and drier ones (Thompson, 1996). Based on the rise and fall of pollen from collected core samples, the mid-Pliocene climate experienced a warmer global climate with enhanced moisture to the western states in part because North Pacific

SST were warmer than today. In conclusion, southern Idaho had overall wetter and cooler summers with warmer and drier winters during the Mid-Pliocene (Forester, 1991;

Ruez, 2006; Thompson, 1996; Thompson and Fleming, 1996).

As mentioned above, Thompson (1996) used a type of cluster analysis

(Constrained Incremental Sums of Squares cluster analysis or CONISS) to define pollen zones that grouped pollen types at different depths of cores (Fig. 15). Zone G-1 through

G-5 represent those groups found in the Glenns Ferry Formation (Thompson, 1996).

Based on the data, there appears to be an oscillation of forest and steppe (plains environment) vegetation during the deposition of the Glenns Ferry Formation

(Thompson, 1996). Initially, zone G-1 (depths 301-281 m) shows a general cooling trend, while zone G-2 (depths 281-252 m) shows a general warming trend and zone G-3

(252-219 m) shows a second cooling trend. However, zone G-4 (depths 219-160 m) and zone G-5 (depths 160-106 m) show a warmer environment near the bottom that trend into a cooler one based on the increase of some pollen taxa (Thompson, 1996).

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Fig. 15. Chart showing the five pollen groups discussed in the reading. I inserted a W (warm) and a C (cool) to indicated the general climate as determined by Thompson (1996) of each row (plant thrives in W or C) and column (overall W or C climate of that pollen sample), (Modified from Thompson, 1996, p. 149.)

Thompson (1996) attributes this pattern to the transgression and regression of a forest boundary relative to a stable lakeshore. However, he does acknowledge another possibility, whereby the forest and steppe vegetation are moving in phase with a lakeshore that is itself moving. However, at zone G-5, Amaranthus and Artemisia increase in correlation with an increase in a warm and dry climate 2.5 to 2.4 Ma

(Thompson, 1996). In addition, this is when the last deep-water phase of Lake Idaho is said to have ceased (Smith, 1987) and the Pleistocene NHG began (Thompson, 1996).

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Fish

Research on Lake Idaho‘s fossil fish was also conducted in order to determine paleoenvironmental conditions. The known temperature limitations (established by oxygen isotope compositions in the calcium carbonate) of related species of fish helped in the reconstruction of temperature variations in Lake Idaho during the Pliocene (Smith and Patterson, 1994).

The presence of both cold and warm water fish species support Lake Idaho‘s similar temperature patterns seen in those of modern fish relatives living in the

Sacramento River Valley today (Smith and Patterson, 1994). Subsequently, the extinction of both the cold water and warm water species also suggests a changing climate regime (a.k.a. the intensification of the Northern Hemispheric Glaciation) near the end of the Pliocene (Smith and Patterson, 1994). The interpretation of the fossil fish shows cool summers and mild winters until about 2 Ma, when lake levels fell (Smith and

Patterson, 1994; Thompson and Fleming, 1996). As stated by Smith and Patterson

(1994), the depletion of 18O in the atmosphere comes from condensation of rain and snow, especially at lower temperatures associated with high elevations and latitudes.

Smith and Patterson (1994) argue that precipitation, from the west-northwest, acted as source water for Lake Idaho that eventually was caught in high altitude snow and ice.

This supports a low 18O found in the carbonates of the fish fossils (Smith and Patterson,

1994).

Using sculpin fish fossils (a small, benthic predator) and supporting fossils like ostracodes from Lake Idaho, Smith (1987) also found stages of a transgressive and

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regressive Lake Idaho based on the lacustrine facies of the Glenns Ferry Formation at

Fossil Creek, Castle Creek, and Birch Creek canyons.

Ostracodes

Ostracodes are microscopic crustaceans (subphylum Crustacea, class Ostracoda)

(Martin and Davis, 2001; Smith and Delorme, 2009). Ostracodes are capable of living in all types of environments including marine, nonmarine, saline, hypersaline, and brackish water (Benson et al., 1961; Pokorny, 1978; Smith and Horne, 2002). They are commonly used as proxies of temperature, salinity, and the general ecology of their habitats because of their sensitivity to change (Benson et el., 1961; Smith and Horne, 2002).

Ostracode distributions are governed by a few forcing factors. These factors include basin morphology, ground-water flow, water depth, water temperature, total salinity, and dissolved ion composition and concentration (Bunbury and Gajewski, 2005;

Forester, 1991; Mourguiart and Carbonel, 1994). Two things to keep in mind while studying ostracodes, however, are their growth through time and the taphonomy of an assemblage. If temperatures are right, individuals can molt up to nine times in their lifetime, producing carapaces and valves that accumulate in the sedimentary record

(Benson et al., 1961; Forester, 1991). This accumulation of identifiable specimens could be disaggregated and transported over time leading to an inaccurate population estimate, by either under- or overestimating the true population present at the time of deposition

(Cronin, 1988; Cronin, 1999). However, if juveniles and adults are present, Cronin

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(1988) agrees that the associated ostracodes were not transported far. Also, depositional environments that include river sands rarely provide fossils due to the ―poor opportunity for preservation and the instability of the sediment‖ (Benson et al., 1961, p. Q57).

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Methods

The ostracodes discussed in this paper were all collected from the Hagerman

Fossil Beds National Monument (HAFO) in the summers of 2007 and 2008. Alison J.

Smith and Donald F. Palmer collected samples HFB-001 through HFB-027 in the summer of 2007. I had the opportunity to stay in Hagerman for the duration of the 2008 summer season, working for HAFO. My field crew and I, collected samples HFB-028 through HFB-075 through the summer of 2008. Locations for the collected samples were constrained by sedimentary composition (explained below), but were distributed as much as possible throughout the park in order to ensure a good overview of the terrain at different elevations.

Field Work

Locations for fossil collection sites were primarily determined by surrounding sediment type. Carbonate-rich clays and silts are usually the ideal sediment for ostracode recovery at HAFO. Sandy sediments do not allow for most ostracode specimens to fossilize properly and often end up destroying specimens that might be present otherwise.

Unfortunately, most of HAFO‘s land surface consists of loose, unconsolidated fine to medium sand. However, erosion and landslide activity have led to the exposure of green clay layers throughout the stratigraphic range of the park. These thin, laminated beds are

46 47

easily recognizable by their color and texture difference. Once the beds were discovered, we dug through the alluvium in outcrop exposures in order to access the best in situ sample possible. In the field, dropping 10% Hydrochloric acid to see if it fizzed confirmed of the presence or absence of carbonates. If the excavated sediment fizzed, we double-bagged approximately a half-gallon sample for further processing in the lab. If the sediment did not fizz in the field we normally did not collect a sample.

In the field, my team and I also used a Trimble GPS system for recording latitude, longitude, and elevation coordinates (above sea level) of each site. These GPS locations were then downloaded into ArcGIS and mapped out with a digital photograph of HAFO at 2-foot intervals. (The finished maps are not available to this research as they are private, park property and unavailable to the public). In doing this work, we subsequently helped to develop and organize all the recorded GPS sites in HAFO collected during the summer of 2008.

Lab Work

The next steps in sample processing of the lab consisted of cleaning samples and picking ostracodes from those cleaned samples. There are many adaptations for successfully cleaning samples in ways that minimize breakage of ostracode carapaces and valves, or otherwise render them useless. We used a modified version of methods described from Colman et al. (1990) and Vance et al. (1997).

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1. A clean and dry Nalgene® beaker was weighed in (gram).

Indurated sediment samples were put into the beaker and

weighed. This was later used to determine the ratio of

valves to grams for each sample.

2. We poured 800 mL of boiling water and 1 tablespoon of

baking soda into the plastic beaker with sediment and

stirred until it was incorporated, covered it and allowed it to

cool.

3. Once the sample had reached room temperature, we added

½ tablespoon of Calgon, stirred, and let stand overnight.

4. For very indurated samples that did not disaggregate after

steps 1-3, we used a second freeze/thaw process.

5. The sample was then frozen and thawed to increase

disaggregation.

6. We used sieving as a means of separating the sediment into

3 subsamples based on grain size. The thawed sample was

wet-sieved through a series of 850-150-100 mesh screens

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into small Whirlpac® bags. All of the sieves were cleaned

ultrasonically after each use.

7. Once the sediment settled in the Whirlpac® bag we

decanted the excess water through a fine mesh of about 60

microns to catch any washouts.

8. The bags were then frozen, usually overnight, before being

freeze-dried again.

9. Once the samples had been freeze-dried, they were ready

for picking.

Almost all of the collected specimens were caught in the 100- micron sieve, which is normally where I started looking for specimens.

Using a microscope at 20X, I would place about 2-5 grams of sediment onto a black, gridded microfossil-picking tray. Using two small (.000) sable paintbrushes (one dry and one wet—dipped in de-ionized water) I picked ostracodes.

All 75 samples were picked, 24 samples yielded valves and carapaces, but eight samples yielded good ostracode specimens suitable for isotopic analysis. From those 8 samples, 26 ostracode samples were

50

submitted for isotope analysis, plus a sample of shell material from lower in the section collected by Rick Forester (1991).

Isotope Analysis

The specimens used for isotopic study were from 8 stratigraphically different localities at HAFO. Usually these valves and carapaces were transparent, sometimes iridescent, and unbroken. Those that were not used for isotopic study were opaque or discolored due to recalcification, or were broken and thus could not be properly identified. However, all of the specimens that could be identified were counted in the total number of specimens per sample.

Specimens collected for 18O and 13C isotopic analyses were sent to Dave

Dettman at the University of Arizona Isotope Lab. The samples at the Arizona Isotope

Lab were run using the procedures described in Dettman et al. (2003), with a precision of

± 0.1 for 18O and ± 0.06 for 13C (1).

―18O and 13C of carbonates were measured using an automated carbonate preparation device (KIEL-III) coupled to a gas- ratio mass spectrometer (Finnigan MAT 252). Powdered samples between 20 and 150 mg were reacted with dehydrated phosphoric acid under vacuum at 70C. The isotope ration measurement is calibrated based on repeated measurements of NBS-19 and NBS-18 and precision is  0.1‰ for 18O and  0.06‰ for 13C (1). Samples were heated under vacuum to 200C prior to measurement.‖ - Dettman et al., 2003, p. 269

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Isotopic values from the shells were relative to the VPDB standard (Coplen et al., 1983). The 18O values were then converted to VSMOW following the equation presented in Coplen et al. (1983), which is written as:

18O VSMOW value = 1.03092 (VPDB) + 30.92

These 18O VSMOW values were then adjusted for the temperature dependence of oxygen isotope fractionation following the formula presented by Kim and O‘Neil

(1997) and discussed in Ito et al. (2003):

3 CaCO3 6 -2 10 ln = 2.78  10 T – 2.89 (0 to 500C) H2O

These values were calculated using a water temperature of 12 C, a reasonable value for shallow, littoral lake settings in mid-latitude environments. Using the Craig

(1961) equation, which defines the global mean water line and the relationship between

δ18O and δ2H, corresponding values for δ2H were then calculated:

δ2H = 8 δ18O + 10 per mil

Statistical Methods

For this project, I used two forms of multivariate statistical analysis on HAFO ostracode data: cluster analysis and principal components analysis. Both analyses are

52

ordination methods, but provide different ways of looking at the data. Principal components analysis (PCA) extracts eigenvectors from a standardized correlation matrix, and these eigenvectors can be graphed as independent axes on which the samples and the variables (species) can be plotted (Davis, 2002). Cluster analysis is a hierarchical agglomerative method that begins with a dissimilarity or similarity matrix. A cluster tree can be developed which shows the relative distances amongst samples or variables

(species) in the dataset (Davis, 2002). Cluster analysis ―joins the most similar observations, then successively connects the next most similar observation to the previous observations‖ (Davis, 2002, p.489). I wanted to find out:

1. How the species (variables) and samples (cases) related to each other?

a. Can independent environmental factors be indentified that would

account for most of the variance in the species dataset?

2. How are the samples related to each other in stratigraphic sequence?

a. Does constrained cluster analysis show facies relationships

identified by the ostracode species occurrences?

Cluster Analysis

I used MVSP Plus (Kovach, 2006) to perform constrained cluster analysis, choosing the Gower similarity coefficient and the farthest neighbor clustering method.

The Gower measures the multivariate distance between samples (or species), with 1

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indicating complete similarity, and 0 indicating complete dissimilarity. After testing other distance measures, I found that the Gower (Davis, 2002) seemed to be able to discriminate among the samples in this dataset better than other distance measures.

I constrained the cluster analysis by keeping the samples in stratigraphic order.

The cluster analysis identified four groups of ostracode assemblages, which can be grouped into different environments (see Results).

Principal Component Analysis

I used MVSP Plus (Kovach, 2006) to perform principal component analysis

(PCA) in order to determine relationships between samples (cases) and variable species, and to determine the independent factors explaining most of the variance in this dataset.

PCA uses eigenvalues and eigenvectors extracted from a correlation matrix

(standardized centered variable x variable correlation matrix). The first few axes

(represented by the eigenvalues) are most important in explaining the variance in the data, and these axes are orthogonal and independent of each other.

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Results

Ostracode Assemblages

Fifteen species of ostracodes from 24 ostracode-bearing samples were identified in this study. There are about 420 species of non-marine ostracodes currently living in North America and many have limited environmental ranges (Delorme, 2001).

These environments can be specifically described (i.e. salinity and temperature ranges) and correlated to that ostracode. However, when multiple species of ostracodes are grouped together slight changes in the environment must be present to accommodate each species (i.e. warm and shallow with the presence of aquatic vegetation and silt-sized sediment). Each assemblage has been described in the following section using Delorme

(1970a, b, c, d, and 1971), Forester (1995), and Forester et al., (2005). The salinity column (represented by only fresh and/or saline) of each assemblage was derived from total dissolved solids (TDS) maps from North American Non-Marine Ostracode Database

(NANODe) (Forester et al., 2005).

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Table 2. Table of HAFO species collected in this study during summers 2007 and 2008.

HAFO Species Table Candona sp. Baird, 1845 Candona crogmaniana Turner, 1894 Fabaeformiscandona rawsoni (Tressler, 1957), Griffiths, 1995 Cyclocypris serena (Koch), 1838 Heterocypris incongruens (Ramdohr, 1808) Ilyocypris bradyi Sars, 1890 Limnocythere ceriotuberosa Delorme, 1967 Limnocythere friabilis (Benson & MacDonald, 1963), Delorme, 1971 Limnocythere itasca Cole, 1949 Limnocythere robusta Delorme, 1967 Endemic Limnocythere sp. A (Genus only) Brady, 1868 Endemic Limnocythere sp. B juvenile (Genus only) Brady, 1868 Physocypria sp. Furtos, 1933 Potamocypris sp. A Daday, 1902 Unknown species

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Fig. 16. Common fossil ostracodes recovered from the Hagerman Fossil Beds NM in this study.

A Fabaeformiscandona rawsoni, left valve, female, HFB 015 at 2813‘, 1 mm scale bar.

B Fabaeformiscandona rawsoni, right valve, male, HFB 015 at 2813‘, 1 mm scale bar.

C Limnocythere friabilis, left valve, female, HFB 030 at 2873‘, 0.5 mm scale bar.

D Ilyocypris cf. bradyi, left valve, female, HFB 028 at 3340‘, 1 mm scale bar.

E Limnocythere ceriotuberosa, left valve, female, HFB 075 at 3342‘, 0.5 mm scale bar.

F Limnocythere ceriotuberosa, left valve, male, HFB 075 at 3342‘, 0.5 mm scale bar.

G Limnocythere robusta, left valve, female, HFB 018 at 2817‘, 0.5 mm scale bar.

H Limnocythere robusta, left valve, male, HFB 018 at 2817‘, 0.5 mm scale bar.

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From this list of specimens, assemblages were organized based on their relative abundance at an elevation at HAFO, and also using constrained cluster analysis.

Depending on the presence or absence of a taxon in HAFO, an environmental reconstruction, or assemblage, can be determined based on established environmental preferences of modern species and estimated environmental preferences of extinct species. The next section discusses these assemblages and describes species autecology

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1) Limnocythere robusta Assemblage

Located at the base of the section, from elevations 2805-2873 ft. at HAFO, the

Limnocythere robusta assemblage consists of the ostracode genera and species listed in

Table 3 and Fig. 17. Based on the presence of F. rawsoni and L. itasca, the assemblage

has a preference for fresh but variable water. This assemblage does not represent a

permanent or consistent lake phase.

Table 3. Limnocythere robusta Assemblage Salinity = Saline (S), Freshwater (F)

GENERA/SPECIES PREFERRED TOTAL RANK SALINITY ENVIRONMENT NO. PRESENT Candona Cosmopolitan 2 7 F-S Candona juv. ― ‖ 619 2 F-S Fabaeformiscandona Both permanent and 83 4 S rawsoni temporary bodies of water. Open water. Fabaeformiscandona ― ‖ 364 3 S rawsoni juv. Limnocythere friabilis Permanent lake in shore 38 6 F juv. zone around .46 m deep. Limnocythere itasca Permanent lake in shore 1 8 F juv. zone around .76 m deep. Limnocythere robusta Only found in Pliocene- 620 1 — juv. Early Pleistocene sediments. Similar to L. ceriotuberosa. Endemic L. sp. B juv. — 49 5 —

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a)

b)

c)

Fig. 17. Graphs and maps of the U. S. from NANODe showing total dissolved solids (TDS) and corresponding maps of modern locations of a) Fabaeformiscandona rawsoni, b) Limnocythere friabilis, and c) Limnocythere itasca (Forester et al., 2005).

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2) Limnocythere friabilis Assemblage

Located from elevations 2873-3046 ft. at HAFO, the Limnocythere friabilis

Assemblage consists of the ostracode genera and species listed in Table 4 and Fig. 18.

Based on the presence of F. rawsoni and L. ceriotuberosa the assemblage represents an oligohaline environment.

Table 4. Limnocythere friabilis Assemblage Salinity = Saline (S), Freshwater (F)

GENERA/SPECIES PREFERRED TOTAL NO. RANK SALINITY ENVIRONMENT PRESENT Candona Cosmopolitan 5 7 F-S Candona juv. ― ‖ 389 1 F-S Fabaeformiscandona Both permanent and 50 5 S rawsoni juv. temporary bodies of water. Open water. Cyclocypris serena Permanent stream around 117 3 F .76 m deep. Limnocythere Permanent lake in shore 1 9 S ceriotuberosa zone around .91 m deep. Limnocythere ― ‖ 1 9 S ceriotuberosa juv. Limnocythere Permanent lake in shore 38 6 F friabilis zone around .46 m deep. Limnocythere ― ‖ 244 2 F friabilis juv. Limnocythere itasca Permanent lake in shore 55 4 F juv. zone around .76 m deep. Unknown juv. — 4 8 —

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a)

b)

Fig. 18. Graphs and maps of the U. S. from NANODe showing total dissolved solids (TDS) and corresponding maps of modern locations of a) Cyclocypris serena and b) Limnocythere ceriotuberosa (Forester et al., 2005).

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3) Candona crogmaniana Assemblage

Located from elevations 3047-3330 ft. at HAFO, the Candona crogmaniana

Assemblage consists of the ostracode genera and species listed in Table 5 and Fig. 19.

Based on the presence of C. crogmaniana, F. rawsoni, I. bradyi, and Physocypria sp., the assemblage has a preference for mostly fresh but variable water. The presence of

Physocypria might indicate a stream input in deeper waters than those of assemblages two or four.

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Table 5. Candona crogmaniana Assemblage Salinity = Saline (S), Freshwater (F)

GENERA/SPECIES PREFERRED TOTAL RANK SALINITY ENVIRONMENT NO. PRESENT Candona Cosmopolitan 8 7 F-S Candona juv. ― ‖ 877 1 F-S Candona crogmaniana Limnetic lake area at 3 11 F substrate-water interface at 10 m deep. Candona crogmaniana ― ‖ 32 3 F juv. Fabaeformiscandona Both permanent and 21 4 S rawsoni temporary bodies of water. Open water. Fabaeformiscandona ― ‖ 837 2 S rawsoni juv. Cyclocypris serena Permanent pond or 5 9 F juv. stream around .3-.76 m depth. Heterocypris Temporary pond or 1 13 F-S incongruens shore zone of around .3 m depth. Ilyocypris bradyi Permanent or 4 10 F intermittent streams around .3 m depth. Ilyocypris bradyi juv. ― ‖ 19 5 F Limnocythere friabilis Permanent lake in shore 2 12 F juv. zone around .46 m deep. Limnocythere itasca Permanent lake in shore 1 13 F zone around .76 m deep. Limnocythere itasca ― ‖ 11 6 F juv. Endemic L. sp. A — 2 12 — Physocypria juv. Permanent pond around 7 8 F .76 m depth. Unknown — 1 13 — Unknown juv. — 2 12 —

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Fig. 19. Graphs and maps of the U. S. from NANODe showing total dissolved solids (TDS) and corresponding maps of modern locations of a) Candona crogmaniana, b) Heterocypris incongruens, c) Ilyocypris gibba, and d) Physocypria globula (Forester et al., 2005).

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a)

b)

c)

d)

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4) Limnocythere ceriotuberosa Assemblage

Located from elevations 3340-3365 ft. at HAFO, the Limnocythere ceriotuberosa

Assemblage consists of the ostracode genera and species listed in Table 6 and Fig. 20.

Based on the presence of L. ceriotuberosa, Physocypria sp., and Potamocypris sp., the assemblage has a preference for an oligohaline environment.

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Table 6. Limnocythere ceriotuberosa Assemblage Salinity = Saline (S), Freshwater (F)

GENERA/SPECIES PREFERED TOTAL NO. RANK SALINITY ENVIRONMENT PRESENT Candona Cosmopolitan 4 7 F-S Candona juv. ― ‖ 456 1 F-S Candona Limnetic lake area at 5 6 F crogmaniana substrate-water interface at 10.4 m deep. Candona ― ‖ 27 4 F crogmaniana juv. Heterocypris Temporary pond or shore 1 10 F-S incongruens zone of around .3 m depth. Ilyocypris bradyi Permanent or intermittent 5 6 F streams around .3 m depth. Ilyocypris bradyi juv. ― ‖ 23 5 F Limnocythere Permanent lake in shore 168 2 S ceriotuberosa juv. zone around .91 m deep. Limnocythere Permanent lake in shore 1 10 F friabilis zone around .46 m deep. Limnocythere ― ‖ 56 3 F friabilis juv. Physocypria juv. Permanent pond around 1 10 F .76 m depth. Potamocypris sp. A Temporary pond or 3 8 F intermittent stream around .3-.6 m depth. Potamocypris sp. A ― ‖ 2 9 F juv. Unknown — 1 10 —

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a)

Fig. 20. Graph and map of the U. S. from NANODe showing total dissolved solids (TDS) and corresponding map of modern locations of a) Potamocypris granulosa (Forester et al., 2005).

Based on NANODe‘s modern interpretation of L. ceriotuberosa’s preferred environmental habitats, the Limnocythere robusta Assemblage has a similar tolerance of environmental parameters with variable temperatures, like hot summers and cold winters, and variable precipitation levels that include oligohaline salinities (Forester et al., 2005).

Based on the present species in HAFO, Lake Idaho was probably not in a permanent or consistent lake stage. However, environmental preferences of associated genera were assumed and used to help represent L. robusta’s environmental preferences in this study.

The next assemblage, the Limnocythere friabilis Assemblage, has a preference for an oligohaline environment based on NANODe‘s interpretation of F. rawsoni and L. ceriotuberosa preferred environmental habitats (Forester et al., 2005).

The next highest assemblage, the Ilyocypris bradyi Assemblage, has a preference for a fresh but variable environment based on NANODe‘s interpretation of C. crogmaniana, F. rawsoni, I. bradyi, and Physocypria sp. (Forester et al., 2005). With the

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presence of Physocypria, it is possible to interpret a stream input coming into the deeper waters of Lake Idaho at this time.

The most recent assemblage, the Limnocythere ceriotuberosa Assemblage, is dominated by a preference for an oligohaline environment based on NANODe‘s interpretation of L. ceriotuberosa, Physocypria, and Potamocypris (Forester et al., 2005).

The presence of Physocypria and Potamocypris indicate a lot of surrounding plant material found in near shore environments with a possible stream input, while high counts of L. ceriotuberosa indicate elevated salinity levels.

Because ostracode distributions in these samples were patchy, only two multivariate statistical analyses were chosen for this study. Further collection of additional samples may produce a more abundant ostracode record that will be suitable for additional statistical analyses.

Principal Components Analysis (PCA) Results

The first four axes of the PCA accounted for 53.7% of my total variance (Table

7). Axis 1 is dominated by Potamocypris, C. crogmaniana, and I. bradyi, which accounts for 17.8% of the variance in the dataset. Axis 2 is dominated by Heterocypris, L. ceriotuberosa, and L. friabilis, which accounts for 13.1% of the variance (Fig. 21). Axis

3 and axis 4 account for 11.9% and 10.9% of the variance, respectively. This means that among all of the specimens I collected from samples at HAFO, Potamocypris, C. crogmaniana, and I. bradyi are less likely to be found present with Heterocypris, L.

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ceriotuberosa, and L. friabilis. The independence of these two assemblages relative to each other in PCA is consistent with their ecological significance, as the former assemblage is indicative of fresh water, and the latter assemblage is tolerant of higher salinities and oligohaline conditions.

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Table 7. Table of eigenvalues and variance of HAFO ostracodes.

Eigenvalues and Variance of PCA Data Table

Eigenvalues Axis 1 Axis 2 Axis 3 Axis 4 Eigenvalues 7.067 4.076 2.746 2.232 Percentage 28.267 16.305 10.984 8.928 Cum. Percentage 28.267 44.572 55.555 64.483

PCA Variable Loadings Axis 1 Axis 2 Axis 3 Axis 4 Candona 0.151 -0.354 -0.074 0.202 Candona juv. 0.293 -0.188 -0.103 0.211 Fabaeformiscandona rawsoni 0.022 -0.09 -0.162 0.123 Fabaeformiscandona rawsoni juv. 0.002 -0.106 -0.164 -0.125 Candona crogmaniana 0.354 0.11 0.03 -0.042 Candona crogmaniana juv. 0.323 0.099 0.025 -0.036 Endemic Limnocythere sp. B juv. 0.022 -0.063 -0.123 0.161 Limoncythere robusta juv. -0.007 -0.001 -0.002 -0.012 Ilyocypris bradyi 0.349 0.101 0.024 -0.031 Ilyocypris bradyi juv. 0.348 0.1 0.026 -0.035 Limnocythere ceriotoberosa 0.011 -0.199 0.521 0.012 Limnocythere ceriotoberosa juv. 0.064 -0.37 -0.09 -0.076 Limnocythere friabilis 0.014 -0.218 0.515 0.006 Limnocythere friabilis juv. 0.036 -0.336 0.438 -0.021 Limnocythere itasca 0.005 -0.15 -0.209 -0.397 Limnocythere itasca juv. 0.022 -0.108 -0.121 0.383 Endemic Limnocythere sp. A 0.013 -0.062 -0.076 0.302 Physocypria 0 0 0 0 Physocypria juv. 0.027 -0.272 -0.226 -0.398 Potomocypris sp. A 0.356 0.114 0.031 -0.045 Potomocypris sp. A juv. 0.357 0.114 0.028 -0.048 unknown 0.333 0.113 0.029 -0.047 unknown juv. 0.025 -0.13 -0.159 0.488 Heterocypris incongruens 0.063 -0.4 -0.161 -0.213 Heterocypris incongruens juv. 0 0 0 0 Cyclocypris serena 0 0 0 0 Cyclocypris serena juv. 0.01 -0.023 -0.023 0.034

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Fig. 21. PCA variable graph with labeled specimens. Based on this graph the upper right group of Potamocypris, C. crogmaniana, and I. bradyi are less likely to be found in groups including L. ceriotuberosa, Heterocypris, and L. friabilis.

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75

Fig. 22. PCA scores graph with labeled HFB-xxx sites. This graph is showing that the two sites HFB-028 and HFB-075 are orthogonal and independent of each other.

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The PCA scores graph shows that two samples are dominating axis 1 and axis 2: samples HFB-075 and HFB-028 (Fig. 22). Sample HFB-075 has a high abundance of L.

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ceriotuberosa, whereas sample HFB-028 is dominated by C. crogmaniana. These samples are independent and orthogonal to each other in PCA, indicating that they are ecological end members in this dataset.

Cluster Analysis

The constrained cluster diagram (Fig. 23) illustrates the similarity between ostracode samples based on the Gower similarity measure and the farthest neighbor clustering method. Four groups are evident in this cluster diagram: 1) elevations 2805 to 2873 feet (HFB-047 to HFB 030C); 2) elevations 2873 feet to 3046 feet (HFB 030B-

HFB 002); 3) elevations 3047 ft to 3330 feet (HFB-010-HFB-031); and 4) elevations

3340 to 3365 feet (HFB-028 to HFB-073). The greatest distances amongst these four groups mark boundaries that are consistent with the ostracode ecology of the assemblages that distinguish fresh water from oligohaline conditions.

Ostracode Zonation and Paleolake Phases

These zones are best illustrated in Fig. 24, which summarizes the results of the cluster analysis and the ostracode paleoecology established in the Ostracode Assemblages section. These four zones, identified as I through IV in valves/gram versus elevation in feet at HAFO, define distinctive ostracode characteristics of the paleo lake. Zones I-IV are also clearly seen in the cluster analysis (Fig. 23). I interpret these four zones as

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Paleolake Phases. Phase I represents the final stage of Lake Idaho at the end of the

Pliocene when the fresh but variable water was dominated by Limnocythere robusta, which became extinct during the Early-Mid Pleistocene. Phase II represents an oligohaline environment with an increase in Limnocythere friabilis, lower precipitation levels and an overall warmer climate. Phase III represents a fresh but variable aquatic environment based on the presence of Candona crogmaniana and its preference for fresh, deep, stable waters. Lastly, Phase IV represents a second oligohaline environment based on the presence of Limnocythere ceriotuberosa and other shallow-dwelling ostracodes like Ilyocypris bradyi, Heterocypris incongruens, and Potamocypris. The PCA supports this interpretation by showing the two end-member assemblages of a freshwater environment and an oligohaline environment as independent of each other.

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Fig. 23. Cluster analysis performed with the Gower General Similarity Coefficient using the farthest neighbor and constrained by elevation. The colored sections represent the four lake stages of Lake Idaho discussed in the text that include the ranges: I) HFB- 047/2805 ft.—HFB-030C/2873 ft. II) HFB-030B/2873 ft.—HFB-002/3046 ft. III) HFB- 010/3047 ft.—HFB-031/3330 ft. IV) HFB-028/3340 ft.—HFB-073/3365 ft.

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81

Fig. 24. Graph of species names and their valve/gram at elevation above sea level. Notice the four lake stages based on the cluster analysis (Fig. 23) and the relative increase in a species for each stage. The valves/gram of a species can show the dominant species of a time and therefore, dictate associated water conditions when that species was dominant.

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Stable Isotopes

For marine and nonmarine environments alike, oxygen-16 and oxygen-18 isotopes are always present in the water. However, the ratio of one oxygen isotope to another (18O/16O) may vary based on the climatic and environmental conditions surrounding that body of water (e.g., precipitation values and evaporation rates). Because

16O atoms are isotopically lighter than 18O atoms, they are preferentially removed from water via evaporation leaving the water enriched in 18O atoms (Reheis and Bright, 2008) when not replenished with more 16O rich water. Ostracodes secrete carbonate shells that gather oxygen from the surrounding water as they are made, and so this ratio can be accounted for within those calcite structures. Generally, ostracode valves with more negative (lower, lighter) values of 18O ratios show periods of decreased evaporation or a steady replacement of the evaporated water from river inlets or tributaries (Reheis and

Bright, 2008). A ―vital effect‖ can be expected, in which there is a metabolic fractionation that occurs. In general, oxygen isotope values from shells can be offset from the water in which they were produced by about 4 per mil (Dettman et al., 1995).

Oxygen isotopic ratios were measured in 26 samples of good-quality (transparent, complete, non-calcified valves or carapaces) ostracodes collected from 8 stratigraphic localities within HAFO in 2008 and 2009, plus a sample collected by Rick Forester

(Forester, 1991) from a lower elevation than those samples collected in the present study.

The results are indicated in the following graphs against elevation (Figs. 25 and 26).

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Fig. 25. Graph between 18O values and elevation (ft.) of ostracode specimens collected from HAFO. Notice the large variation low in the section from heavy to light to heavier to heavy 18O values.

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Fig. 26. Graphs showing the relative 18O and 13C values for each of the collected HAFO species. Notice that L. robusta does not occur again after 2840 ft.

Oxygen isotope values range from -2 to -18 per mil VPDB in the ostracode shells from HAFO. The lightest isotopic values are measured in Limnocythere friabilis, a species associated with groundwater discharge, which produced a 18O value of -18 per mil VPDB. Cytherissa lacustris, a benthic, open lake species that prefers permanent cold water with low salinity, was collected in a sample by Rick Forester (Forester, 1991) from a section just below the lowest elevation sampled in this study, and produced a value of -

17 per mil VPDB. The water in which it lived had light isotopic values of at most -17 per

86

mil VSMOW, which is at the lower isotopic limit of summer atmospheric precipitation values in Idaho today (Benjamin et al., 2004). Most of the other species range in 18O values between -3 and -12 per mil VPDB, which correspond to atmospheric precipitation

18O values of -3 to -13 per mil VSMOW. A charophyte sample (P in Fig. 26) taken in the upper section at HAFO (a gyrogonite from a species of charophyte) also produced light oxygen isotope values of about -15 per mil VPDB. Carbon isotopes are harder to interpret in this context, because ostracode carbon isotope signatures are possibly a mix of the total inorganic carbon and the carbon they consume (Dettman et al., 1995).

If the VPDB values of 18O from HAFO are converted to VSMOW and adjusted for temperature (see methods section), they can be used to calculate a corresponding value of 2H. These values, then, the temperature adjusted 18O VSMOW values calculated from VPDB, and the calculated 2H values can be plotted on the modern local mean water line for southeastern Idaho (Fig. 27) (Benjamin et al., 2004). What is important to note about Fig. 27, is that the HAFO values plotted on the modern water line fall within the normal range of summer and winter precipitation levels seen in Idaho today. Therefore, the source for precipitation in Idaho during the late Pliocene – early

Pleistocene must have been very similar to the present day source, which supports modern seasonality and ENSO cycles.

This general approach sheds light on the seasonal or non-seasonal source of precipitation. Ultimately, given more isotope data from additional material, an evaporation line for the paleolake could be calculated.

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Fig. 27. Modified image of the modern local meteoric water line for southeastern Idaho, western Wyoming, and south-central Montana. The diamond HAFO values fall along the water line, which represents the modern seasonality cycles of summer and winter precipitation. (Modified from Benjamin et al., 2004, Figure 4).

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Discussion

Hagerman Fossil Beds National Monument appears to have experienced four primary lake stages that transitioned between a fresh but variable environment to an oligohaline environment throughout the late Pliocene – early Pleistocene time. During the first fresh but variable stage, HAFO appears to have received more frequent precipitation based on the 18O isotopes of HAFO ostracodes (Fig. 17), the Limnocythere robusta Assemblage (Table 3), and supporting pollen data, which shows higher levels of

Pinus associated with cooler, wetter conditions (Thompson, 1996). At the time of deposition, higher North Pacific sea surface temperatures helped to increase precipitation to areas that now make up the arid western states (Bartoli et al., 2005).

However, based on the Limnocythere robusta Assemblage (Table 3), the lake was not a stable, consistent, permanent lake. Compared to a similar assemblage found in

Beaver Basin, UT, L. robusta does not reappear after 2 Ma (Forester and Bradbury,

1981). As seen with a lack of L. robusta in any later HAFO assemblage, the

Limnocythere robusta Assemblage is thought to represent the final days of Lake Idaho during the late Pliocene.

The next lake stage shows lower annual precipitation levels and warmer temperatures. An oligohaline environment has been identified based on the 18O isotopes

(Fig. 18) of the HAFO ostracodes and the associated Limnocythere friabilis assemblage

(Table 4). Based on pollen data collected from the surrounding area, the regional area

88 89

was experiencing a warmer, drier climate with elevated counts of Ambrosia-type vegetation. This supports an oligohaline lake because of the associated lower precipitation rates and higher associated evaporation rates of warmer climates.

The third lake stage indicates a second fresh but variable lake environment.

Based on the pollen record, there was another peak in Pinus associated with cooler, wetter conditions (Thompson, 1996). The Candona crogmaniana Assemblage (Table 5) is also conducive to a fresher, deeper, more stable lake associated with higher precipitation levels relative to evaporation rates. Lake Idaho most likely had a stream input, which provided a consistent water supply along with seasonal precipitation. Also,

O18 levels also show a more fresh water environment than stage one (Fig. 19). This stage is thought to be the primary deep-water stage of Lake Idaho‘s lifespan.

Lastly, the fourth lake stage is indicative of an oligohaline environment with elevated salinity levels and associated warmer, drier conditions. Based on pollen data, the fourth stage starts off warm then trends into a cooler climate associated with an another climatic control called the Pleistocene Northern Hemisphere Glaciation (NHG), which started approximately the same time Lake Idaho experienced its last deep-water phase 2.5-2.4 Ma (Thompson, 1996). Notice also that the O18 levels show a final trend towards more saline environments associated with those of desiccating bodies of water

(Fig. 20).

In other studies, there are similar findings between seasonal variations in the

HAFO area (Forester, 1991; Thompson, 1996) and other western North American areas of the same time period (Forester and Bradbury, 1981). With that in mind, all the

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ostracode isotopic records discussed above, as well as the pollen records, from HAFO indicate that a modern seasonal ENSO, with fluctuations between a La Nina and El Niño event, were already in place during the deposition of HAFO after the presence of Lake

Idaho and into the early Pleistocene (see Results).

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Conclusions

Lake Idaho, based on an excess of precipitation relative to evaporation, was a large and permanent Pliocene lake, located in the southwest portion of Idaho and parts of eastern Oregon (Fig. 28) (Smith and Patterson, 1994; Ruez, 2006). Located within the

Glenns Ferry Formation of the Idaho Group, deposition of Glenns Ferry Lake began after the deep-water phase of Lake Idaho, indicated in this project by the last appearance of

Limnocythere robusta. Glenns Ferry Lake existed as a shallow lake until its infilling after deposition of the Fossil Gulch Ash and after deposition of the Hagerman Horse

Quarry. Thompson (1996) and Kimmel (1982) concluded that Lake Idaho had an outlet based on the absence of chemical precipitates within sediment samples and lack of evaporite deposits. This outlet was most likely located towards the Columbia River drainage system, where erosion may have ultimately caused drainage of Lake Idaho around 2.4 Ma (Kimmel, 1982; Thompson, 1996).

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Fig. 28. Hypothetical location of Lake Idaho based on research conducted on fish in the surrounding area (Smith and Patterson, 1994, p. 293).

Using ostracodes as paleoenvironment indicators I have shown that during the latest Pliocene and mostly of the early Pleistocene, Lake Idaho experienced at least four lake stages before completely collapsing.

The first stage was that of a fresh but variable aquatic environment with supporting levels of Pinus indicative of a cool and wet climate. The second lake stage is indicative of an oligohaline environment with warmer, drier conditions. We also found

93

higher levels of Ambrosia-type vegetation (a kind of ragweed) that prefer warm, dry climates. The third lake stage shows a second fresh but variable environment with cooler, wetter conditions than lake stage two and fresher water than lake stage one. This is also associated with a peak in Pinus, associated with cooler, wetter climates. The fourth, and last, lake stage is that of an oligohaline environment with slightly elevated salinity levels than that of lake-stage two.

HAFO‘s depositional environment was affected by modern northwest-west precipitation sources. Precipitation sources consistent with these modern sources indicate that a seasonal ENSO cycle was already in place. Based on pollen data, lake stage four trended into a cooler climate, which may have been the beginning of the Northern

Hemisphere Glaciation (NHG) at the end of the Pliocene into the early Pleistocene.

These four aquatic stages are identified by ostracode species data. The cluster analysis further emphasizes their stages. Based on the early presence and immediate disappearance of L. robusta in later stages, we have concluded that most of HAFO was deposited during the late Pliocene and earliest Pleistocene. The oxygen isotope analysis of well-preserved shells further indicates that a seasonal ENSO was already in place at the time of deposition. Pollen and fish proxies analyzed by other works (Smith and

Patterson, 1994; Thompson, 1996) from the immediate and surrounding area have also helped in the determination of environmental parameters.

Suggested further research includes the continued study of ostracodes and limiting depositional beds in Hagerman Fossil Beds National Monument, Idaho as well as the surrounding area. These studies should focus on the isotopic record and associated

94

paleoenvironmental limitations of the fossils and consider addressing a more restrictive time constraint for the park. Hagerman Fossil Beds National Monument is an abundant resource of Plio-Pleistocene paleoecological transitions and signals of global climatic shifts. As the study of ostracodes evolves, HAFO‘s high desert beckons with its trove of discoveries to come.

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