A Sedimentary Record of Holocene Neoglaciation from the Uinta Mountains, Utah

A SEDIMENTARY RECORD OF HOLOCENE NEOGLACIATION FROM THE UINTA MOUNTAINS, UTAH

Catherine Klem

Submitted in partial fulfillment of the requirements for the degree of Bachelor of Arts Department of Geology Middlebury College Middlebury, Vermont

May 2010 Klem, Catherine M. 2010. A Sedimentary Record of Holocene Neoglaciation from the Uinta Mountains, Utah. Unpublished Senior Thesis: Middlebury College, Middlebury Vermont. 77 pgs.

ABSTRACT

Considerable paleoclimate research is focused on climatic variability during the Holocene, including the enigmatic period of renewed glaciations known as the Neoglaciation. Alpine lake sedimentary records are particularly useful in Neoglacial studies because they can provide uninterrupted, high-resolution records of environmental changes. This study focused on analysis of biogenic silica, phosphorus, and carbon:nitrogen from an alpine lake sediment record collected in northeastern Utah, adding to an extensive dataset that suggests several periods of Neoglacial activity in the Uinta Mountains.

The sediment core, which spans the past 5300 years, was taken from EJOD Lake, a small lake with a maximum depth of 4.3 m, and a surface area of 2.7 ha, located above modern treeline at 3323 m. The lake is situated in a large glacial cirque ~500 m downslope from a complex of Neoglacial end moraines. Dry meltwater channels run from these moraines directly into EJOD Lake, thus establishing a clear connection between glacial activity and the lake sediment record.

Biogenic silica, loss-on-ignition, and clastic flux time series all exhibit an overall decreasing trend over the past 5000 years, suggesting a gradual shift of the lake environment towards a colder, less productive climate. Superimposed on this downward trend are smaller-scale fluctuations characterized by intervals of notably low values in biogenic silica and LOI, and high values in C:N. These intervals suggest episodes of Neoglacial advance at 3600-3400, 2500-2300, and 2000-1600 BP. Subsequent increases in mineral P and clastic flux and decreases in median grain size at 3000-2900, 2300-2000, and 1400-1200 BP mark times when glacial retreat washed fine-grained rock flour out of the moraines into the lake. Less pronounced variability at 800-550 and 400-200 BP show similar trends which may indicate two episodes of Neoglacial advance associated with the Little Ice Age. The dating of these episodes shows some synchronicity with periods of Neoglacial advance in western North America, but also indicates that such advances were extremely regionally variable. A dramatic shift in proxies indicates a decline in soil development between AD 1750 and AD 2005 likely due to the onset of grazing ca. AD 1850 in the area.

i ACKNOWLEDGEMENTS

First and foremost, I would like to thank Jeff Munroe, my thesis and academic advisor. You patiently answered all of my questions, introduced me to lab work, and somehow always made me believe that this was possible. You taught me everything from how to use a repeat pipette to how to efficiently navigate excel, and I really appreciate that.

Also, thank you to the entire geology department, both the professors and my fellow seniors. I can confidently say to all of my peers that I studied in the coolest department and whether my future studies delve deeper into geology or lead me to other endeavors, I will always greatly cherish my Middlebury education.

And to my friends and family who have supported me throughout this project and who listened to my stories of long hours in bihall and entertained me while in the lab. Thank you for the moments when you offered a little perspective on life and for the moments when you supported my inner-geek.

Finally, thank you to the Middlebury College Senior Work Fund for supporting my research and to the Middlebury College travel fund for making it possible for me to travel to Maryland and see what a real geology conference is all about.

ii TABLE OF CONTENTS Abstract...... i Acknowledgements…………………………………………………………………….….ii Table of Contents ………………………………………..……………………...………..iii List of Figures and Tables …………………………………………..……...……..……..iv Chapter I: Introduction………………………………...... ……..1 The Importance of Paleoclimate studies…………………………………………..1 Holocene Neoglaciation……………………………..…………………………….2 The Study of Paleoclimate……………………………………..………………….7 Climate Proxies……………………………………..……………………………10 Chapter II: Study Site…………………………………………………………………….13 Uinta Mountains………………………...………………….……………….……13 EJOD Lake……………………………………………....……………………….15 Previous studies of Glacial History in the Rocky Mountains…………………....20 Chapter III: Methods…………………………………………………………….……….22 Field Work……………………………………………………………………….22 Previous Work on EJOD Lake Core………………………………….………….23 Sample Analysis………………………………………………………….………23 Chapter IV: Results………………………………………………………………………27 Age Control…………………………………………………………………..…..27 Detrending………………………………………………………………………..28 Biogenic Silica……………………………...... ……..…….28 Phosphorus……………………………...... ….……………31 C:N……………………………...... …………………………….34 Flux………………………………………………………………………..……..36 Overall Time Series………………………………………………………...……39 Chapter V: Discussion……………………………………………………………...……43 Relationships between Proxies…………………………………………………..43 Interpretation of EJOD Lake History………………………………….……..…..51 Links to Regional and Global Climate………………………….………..………59 Chapter VI: Conclusions…………………………………………………...... ………….69 Sources Cited…………………………………………………………………………….71

iii LIST OF FIGURES AND TABLES

Chapter I: Figure 1.1 Intervals of Holocene glacial activity in the Western U.S.…………………....5

Chapter II: Figure 2.1 Digital Elevation Model of the Uinta Mountains, UT………………………..16 Figure 2.2 Bathymetric map of EJOD Lake…………………………………………..…16 Figure 2.3 Topographic map of the West Fork of the Black’s Fork……………………..17 Figure 2.4 Photo of the landscape surrounding EJOD Lake…………………….…….…17 Figure 2.5 Photo of Meltwater Channels………………………………………………...18 Figure 2.6 Image of the Surrounding Topography……………………………………....19

Chapter III: Figure 3.1 Photo of lake coring operation……..………………………...... ……..22 Figure 3.2 Method for differentiating biologically and minerally derived silica………..24

Chapter IV: Table 4.1 AMS ages used for depth-age model…………………………………….……27 Figure 4.1 Depth-age model……………………………………………………………..27 Figure 4.2 Biogenic Silica time series………………………………………...…………30 Figure 4.3 Total P time series……………………………..……………………………..32 Figure 4.4 Mineral P time series…………………………………………………………33 Figure 4.5 C:N time series……………………………………………………………….35 Figure 4.6 Flux times series……………………………………………………………...37 Table 4.2 Summary statistics………………………………………………………….....41 Figure 4.7 Summary plot………………………………………………………………...42

Chapter V: Figure 5.1 Summary plot with additional proxies…………………………………...…..43 Figure 5.2 Comparison of color data with % mineral P and clastic flux……………...…51 Figure 5.3 Episodes of Neoglacial advance in EJOD Lake record………………………53 Figure 5.4 Comparison of EJOD Lake proxies……………………………….………….54 Figure 5.5 Summer insolation data………………………………………………………61 Figure 5.6 Comparison plot of Neoglacial advance in the Western U.S……………..….64

iv I. INTRODUCTION

The Importance of Paleoclimate studies

The study of climate change is an important field that is now more relevant than ever. Research has shown that between AD 1860 and present day the concentration of

CO2 gas in the Earths atmosphere has increased from 280 to 387 ppmv contributing to an increase of 0.55°C in the global mean annual surface temperature (Lean et al., 1995;

NOAA, 2010). Such changes have been largely credited to anthropogenic influences, and could have widespread implications for the future, affecting both plant and animal habitat. However, predicting and understanding human-driven climate changes is greatly complicated by the presence of natural cycles and variability (Lean et al. 1995).

Accordingly, a good understanding of past climate cycles and variations is necessary to provide a basis with which to evaluate current climate observations (Kaplan et al., 2002).

One of the best ways to better understand these natural variations, and consequently the interface between natural and anthropogenic factors, is by considering past cycles and changes through the study of paleoclimate (Laird et al., 1996; IPCC report, 2007).

Paleoclimate reconstructions can shed light on the scale and rapidity of climatic and environmental changes we face today as well as help investigate the forces that drive climate shifts and the earths reactions to such changes. Recently, studies of Holocene

Neoglaciation (a period of locally renewed glaciation after the last Pleistocene Ice Age) have become important for understanding paleoclimate drivers and the global synchronicity of periods of warming and cooling (Schaefer et al., 2009). Holocene

Neoglaciation studies are especially important as they are on a time scale that humans can

1 relate to and even play a part in, and thus show an interesting interface between natural and anthropogenic climate shifts.

The primary goal of this thesis is to carry out a multi-proxy study of Holocene

Neoglaciation from a high alpine lake sediment record (EJOD Lake) in the Uinta

Mountains of northeastern Utah (Fig. 7). Newly analyzed proxies will be used in conjunction with existing data to reconstruct climate variability during Holocene

Neoglaciation over the past 5300 years, and will be compared to other climate reconstructions from the Rocky Mountains. A more complete understanding of climate records for the study site and their relation to the surrounding environment will help bolster the paleoclimate record for the Rocky Mountain region which can in turn be used to better understand present and future climate implications.

Holocene Neoglaciation

The Holocene is a geologic epoch in the Quaternary Period which began approximately 11,500 years ago and continues to the present day. The epoch is marked by an interglacial warm climate, which has provided a hospitable environment, marked by mild temperatures and precipitation (Beget, 1983). While until recently the Holocene was believed to be a relatively stable period, especially when compared to the notably varied climate of the Late Pleistocene, it is now understood to have undergone definite climatic variability and abrupt changes (Rietti-Shati et al., 1998; deMenocal et al., 2000).

Schaefer et al. (2009) suggest that the Holocene epoch has been marked by a dominant cooling trend underlying millennial-scale variations. The early to middle

Holocene (8,000 to 3,000 yr B.P.) was marked by a period known as the Hypsithermal, a warm interval which resulted due to a combination of Milankovitch cycles and the

2 disappearance of the North American ice sheet (Whitlock and Bartlein, 1993).

Palynological studies have estimated that temperatures during this period were comparable to, if not warmer than present day conditions. Consequently, during this period, ice, which had been widespread during the glacial period, vanished from many settings (Whitlock and Bartlein, 1993).

The Hypsithermal warm interval was followed by a period of Neoglaciation during which much of the ice which had melted during the Hypsithermal returned. The

Neoglacial period was marked by several thousand years of climate fluctuations, including several periods of glacial advance, especially in alpine environments (Porter,

1986; Beget, 1983). The onset of Neoglaciation is generally defined by an initial episode of glacial advance around 5000 BP (Porter, 2000). A study by Beget (1983) however used radiocarbon dating in conjunction with glacial deposits from several regions of the world to indicate that perhaps Neoglaciation began as early as 8500-7500 BP. A more recent study by Douglass et al. (2005) used cosmogenic nuclide surface-exposure dating to show that indeed glaciers in South America may have begun to advance about 3000 years earlier than previously thought, with advances at 8500 and 6200 BP. Nevertheless, climate variability was perhaps the greatest in the last 3000 years, culminating in the

Little Ice Age, when most Northern Hemisphere glaciers reached their maximum

Holocene extent. The Little Ice Age was then followed by a period of significant warming (Kaplan et al., 2002).

While it is quite difficult to generalize the more recent periods of glacial advance that mark Holocene Neoglaciation due to the high amount of regional variation, studies investigating these fluctuations throughout the Rocky Mountains have found some

3 dominant trends (Fig. 1.1). A study by Fall (1997) used pollen data from sediment cores to better understand climate fluctuations since 9000 BP and found that between 9000 and

4000 BP the climate was 1-2ºC warmer than current conditions, with conditions becoming progressively drier over that time period. Munroe (2003) similarly used pollen data and found that the Rocky Mountains reached their maximum Holocene temperatures between 6500 BP and 5400 BP and since then conditions have been very similar to those of present day. Smaller scale fluctuations varied greatly throughout the Rocky

Mountains but there is evidence of glacial advance in the Front Range of Colorado

(Benedict, 1993) and the Wind River Range of Wyoming (Dahms, 2002) between 5000 and 3500 BP. Lichenometry of rock glaciers in the Colorado Front Range show that these features advanced in 2070 BP and 1150 BP (Refsnider and Bruffer, 2007). Records from the Uinta Mountains (Munroe, 2002), the Sawatch Mountains (Refsnider and

Brugger, 2007), the Front Range (Benedict, 1993) and the La Sal Mountains (Nicholas and Butler, 1996) also indicate another period of glacial advance between 2400 BP and

2000 BP. Most recently, throughout much of the Rocky Mountains there was a period of glacial advance, known as the Little Ice Age, between AD 1350 and AD 1860 (Konrad and Clark, 1998). Similarly, Clark and Gillespie (1996) used cirque moraines in conjunction with lake sediment records to show that pronounced glacier advances occurred in the Sierra Nevada around 650 BP (Clark and Gillespie, 1996).

In the United States, the earliest geomorphic evidence of Neoglacial fluctuations can be found on Dome Peak in Washington. Radiocarbon analysis of tree trunks, sheared by the advancing glacier, dates the deposition of end moraines there to between 5000 and

4600 BP (Davis, 2009). Evidence of the Middle Neoglacial period is best seen in moraines in the Rocky Mountains of Colorado and Wyoming, which were deposited between about 2000 and 1000 BP (Davis, 2009). Moraines from the Little Ice Age cooling event dominate the alpine glacial landscape throughout North America due to the fact that in most cases glaciers reached their maximum extent during the LIA, thus

5 destroying most evidence of previous events. These moraines can be identified by their sharp-crests and unweathered appearance (Davis, 2009).

Despite geomorphic evidence and extensive research, many questions remain regarding Holocene Neoglaciation. For one, the reason for glacial advance is still not well understood. Wanner et al. (2008) suggests that the primary climate drivers for the last 6,000 years were variations in orbital position, solar forcing and volcanic forcing, where orbital position describes the position and orientation of the earth relative to the sun; solar forcing relates to the varying amounts of radiation emitted from the sun; and volcanic forcing is caused by the emission of sulfur gases during volcanic eruptions which causes cooling of the atmosphere long after the eruption has stopped (Wanner et al., 2008). Wanner et al. (2008) also suggested that major land cover changes and changes in greenhouse gases may have played a secondary role. Research by Holzhauser et al. (2005) supports the theory that the main driver has indeed been variation in solar radiation. Yi et al. (2008), however argue that Neoglaciation has in fact not been caused by Milankovitch cycles but rather reflects more local changes in precipitation that were tied to the general cooling trend of the Holocene.

Another major question regarding Holocene Neoglaciation is the extent to which glacial advances have been globally synchronous. In a study of maximum glacial extent,

Davis (2009) concluded that Neoglaciation is not globally uniform based on the findings that glaciers in the Colorado Front Range have become successively smaller throughout the Holocene, while glaciers in the rest of North America have increased in extent, and thus concluded that Neoglaciation is not globally uniform. A more recent study by

Benson (2007) found similar results, concluding that there is no indication of global or

6 even North American synchronicity. Domack and Mayewski (1999) investigated the same question by comparing marine sediment cores from the Antarctic Sea and ice cores from Greenland. No conclusions were made as to which hemisphere forced the other, but

Domack and Mayewski (1999) concluded that there are definite links between climate change in the northern and southern hemisphere. In a very recent study, Schaefer et al.

(2009) compared climate records from glaciers on Mount Cook in New Zealand to studies from North America, drawing the conclusion that the maximum ice extent occurred at significantly different times between the northern and southern hemispheres.

In general Mount Cook glaciers reached their maximum extent around 6500 years ago while northern hemisphere glaciers did not reach their maximum extents until AD 1300 to AD 1860. Schaefer et al (2009) did, however, find that the greatest coherency between the Mount Cook and Northern Hemisphere records occurred between AD 300 and AD

700 with some consistency in the past 700 years as well.

The Study of Paleoclimate

Several different approaches have been used to study paleoclimate including glacial reconstructions, lake sediment cores, ice cores, packrat middens, pollen assemblages, plant macrofossils, and dendro-climatology (Refsnider and Brugger, 2007).

In the Rocky Mountains, due to the widespread presence of glacial deposits and past alpine glacier activity, studies have traditionally focused on the presence of geomorphic landforms created by glaciation to reconstruct and date past glacial extents. Moraines have been dated using Cosmogenic Surface-Exposure Dating, a method by which the length of time a boulder has been exposed at the surface is calculated using established production rates of radioactive nuclides (Douglass et al., 2005; Munroe et al., 2006).

These dating techniques have been used in conjunction with glacial modeling, relying on the principle that glacier size is directly related to climate conditions and that fluctuations in glacier extent can be used as a proxy for paleoclimate (Laabs et al., 2006). Further, glacial extent has been calculated based on the position of lateral moraines and breaks in the slope on cirque headwalls (Laabs et al., 2006).

One major issue with the use of these methods in North America is that the Little

Ice Age, the most recent cooling event, produced in many places the most extensive and far reaching glaciers. These glaciers wiped out older moraines and other topographical evidence. The result is that in many places only records of the most recent advances are present in the landscape (Thackray, 2008).

One method that is a good indicator of the frequency and magnitude of climate change, and does not rely on the presence of geomorphic features, is the study of lacustrine sedimentary records from middle-latitude, high alpine lakes (Laabs and

Carson, 2005; Matthews et al., 2005). The high-resolution record, frequent presence of annual lamination and dateable fossils, and the direct link to groundwater and precipitation in the surrounding basins, all make lake deposits ideal paleoclimate indicators (Pederson, 2000).

Sedimentary records of alpine glacier fluctuations are particularly useful, because, due to their relatively small size, alpine glaciers are very sensitive to fluctuations in the climate and thus provide a very detailed history of past glacial advance and retreat in their sedimentary record (Abbott, 1997; Refsnider et al., 2008; Schaefer et al., 2009;

Willemse and Törnqvist, 2009). Porter (1986) explains that large and cold high-latitude glaciers have longer response times, but small low and middle-latitude glaciers are very

8 sensitive to climate change due to a very short lag in the response time of the glaciers to environmental changes. This high level of sensitivity is especially important in the study of Holocene Neoglaciation as fluctuations were generally of high-frequency and low magnitude (Porter, 1986). Alpine glaciers are also responsive to climate change on both regional and global scales. Not only do their fluctuations reflect regional changes in precipitation and temperature, but they are also influenced by changes in atmospheric circulation over the Pacific Ocean and North America, and thus correlate to changes in the Southern Oscillation Index and Northern Hemisphere air temperature, providing important insight on global climate cycles (Laabs et al., 2006).

In the analysis of sediment lake cores, there are several commonly applied proxies including grain size, loss-on-ignition (LOI), magnetic susceptibility, bulk density, biogenic silica, carbon:nitrogen, and mineral phosphorus, which can give insight on temperature, precipitation and other important lake properties (Matthews et al., 2005).

Grain size analysis on lake cores can show variations in the particle size of sediments deposited in the lake as glacial runoff or from other sources. Rodbell et al. (2008) investigated the viability of grain size as a proxy for glaciation during the Late

Pleistocene and found that periods of increased clastic input indeed aligned with ages of cosmogenically dated erratics on moraine crests (Rodbell et al., 2008). Loss-on-ignition analysis is used as an indicator of glacigenic sediment content where low LOI values indicate the presence of glaciers in the watershed. A lake sediment study done in

Greenland used a combination of LOI and biogenic silica (bSi) analysis to identify cold periods in the sediment record and define periods of Neoglacial advance (Kaplan et al.,

2002). Magnetic susceptibility varies directly with minerogenic content of sediments and

9 thus is an indicator of glacial activity. In a study in northern Sweden, mineral magnetic measurements shed light on the intensity of the spring snow melt, and concentration of organic carbon (Snowball et al., 1999). Water content and bulk density are inversely related to glacier activity (Matthews et al., 2005). Biogenic silica contents are used to quantify lake-water properties and air and water temperatures (Bigler et al., 2002). A study by Rietti-Shati (1998) used biogenic silica concentrations in a sediment record from a high alpine lake on Mount Kenya, Africa to reconstruct temperature data between 4200 and 1200 BP (Rietti-Shati, 1998). Carbon:nitrogen comparisons can help determine whether organic matter in sediments is of terrestrial or aquatic origin (Qiu et al., 1993), and relative percentages of mineral phosphorus can be used as a measure of mineral erosion (Filipelli et al., 2005). A study in southwestern Alaska used a combination of bSi concentrations and C:N ratios to suggest periods of increased N cycling in the lake which was reflected in changes in Alnus pollen concentrations (Hu et al., 2001).

Climate Proxies

In this study, three climate proxies will be used to better understand climate fluctuations and changing lake properties over time. Biogenic silica (bSi) analysis will be used to determine biogenic silica abundance, a proxy for in-lake productivity. Biogenic silica is derived from the siliceous skeletons of diatoms, which are relatively ubiquitous microscopic algae in freshwater lakes (Moschen et al., 2006; Korhola and Weckstrom,

2000). When the diatoms die, their siliceous skeletons are deposited on the lake bottom and thus preserved in the sediment record (DeMaster, 1981). Because diatoms are dominant primary producers in most lakes, bSi is a good indicator of lake productivity

(Hu et al., 2001). Furthermore, Kaplan et al. (2002) suggests that because bSi directly

10 measures in-lake algae, it is a good proxy of aquatic activity. Productivity and biogenic silica abundance have been found to be influenced by several factors including air temperature, water temperature, nutrient cycling, lake level, and lake-water pH (Bigler et al., 2002; Kaplan et al., 2002; Moschen et al., 2006; Qiu et al., 2009). Studies show, however, that the presence of diatoms (and consequently biogenic silica) is most significantly affected by fluctuations in mean summer air temperature (lower water temperatures form inhospitable environments for the diatoms) and landscape erosion

(increased sediments eroded off the landscape increases lake turbidity which in turn blocks sunlight necessary for diatom photosynthesis) (Korhola and Weckstrom, 2000).

Decreased biogenic silica concentrations in the sediment record consequently suggest cold periods dominated by landscape erosion, while increased abundance suggests warmer periods.

Carbon:nitrogen analysis is used in this study to help to identify the organic matter source (terrestrial vs. aquatic) during sediment deposition. The carbon:nitrogen technique is based on the premise that, due to their thick cell walls, terrestrial plants are rich in carbon, while the much less rigid aquatic organic material is richer in nitrogen

(Sampei and Matsumoto, 2001). Consequently a ratio of carbon to nitrogen can be used to assess whether deposited organic matter is dominantly of terrestrial or aquatic origin

(Qiu et al., 2009). C:N analysis is especially helpful when used in conjunction with LOI data to determine the nature of LOI changes. LOI data from a previous study provide valuable information as to the amount of organic matter being deposited on the lake bottom, but do not specify the origin of the organic matter. Consequently a combined

C:N and LOI analysis will help determine both the amount of organic matter present at a

11 given time, and whether or not it is aquatically derived (and therefore representative of lake productivity).

Finally, mineral phosphorus analysis will be used as a measure of glacial erosion.

The relative proportions of three pools of phosphorus mineral, occluded, and organic will be compared. While the total phosphorus in a lake varies based on numerous different factors, the relative abundance of mineral, occluded and organic phosphorus can help determine sediment origin (Filipelli et al., 2006). Over time as soils develop, mineral phosphorus is converted to occluded and organic phosphorus, where occluded phosphorus is that which has been co-precipitated with or adsorbed onto oxyhydroxides

(most notably iron oxyhydroxides). Thus increased percentages of occluded and organic phosphorus signify increased soil development. Contrarily, because the original source of mineral phosphorus is the bedrock, a rise in mineral phosphorus signifies increased bedrock erosion, which in an alpine lake environment in close proximity to a cirque glacier would likely have been caused by glacial advance (Filipelli and Souch, 1999;

Filipelli et al., 2006). This analysis is best done in oligotrophic lakes where primary production is low and sediment records reflect landscape erosion (Filipelli et al. 2006).

Biogenic silica, C:N and phosphorus analysis will be combined with color, LOI, grain size, and bulk density data previously collected by J. Munroe and considered collectively to identify trends in the data that may signify periods of warming or cooling

(J. Munroe, unpublished data). Once these trends have been identified, their timing will be compared with other records of Holocene Neoglacial advances elsewhere in the Rocky

Mountains.

II. STUDY SITE

Uinta Mountains

A subset of the Rocky Mountains, the Uinta Mountains run east-west for 200 kilometers along the Utah-Wyoming border, and are located downwind of the paleoshorelines of the 51000 km2 pluvial Lake Bonneville (Thackray, 2008). The highest mountains in Utah, the Uintas range in height from 3400 to 4100 meters with the highest peak being Kings Peak at an elevation of 4136 meters above sea level (Thackray, 2008).

The peaks can be divided between the western peaks, which have been glaciated, and the eastern peaks, which have not (Munroe, 2005).

The Uinta Mountains are particularly well suited to lake sediment studies of

Holocene Neoglaciation due to the presence of widespread Neoglacial deposits even in places where glaciers no longer exist today (Laabs et al., 2006). Further, well-preserved moraine ridges from at least the past two glaciations provide important supplementary information (Laabs and Carson, 2005). Also of note is the Uinta Mountains unique east- west trending orientation which runs parallel to storm tracks and moisture transport thus facilitating the study of paleoprecipitation gradients (Munroe et al. 2005). Paleoclimate records from the Uinta Mountains also serve as an important link between records from the middle Rockies, Colorado Plateau, and the Great Basin/Sierra Nevada (Munroe et al.

2005).

The Uinta Mountains are predominantly made up of orthoquartzite, shale, and sandstone units from the Uinta Mountain Group (UMG), a 7 km thick sequence that was deposited around 850-750 Ma in an intracratonic basin (Munroe, 2005; Dehler and

Sprinkel, 2005). In some places, a 50-m thick sandstone layer from the Cambrian Tintic

and Lodore formations overlies the Red Pine Shale (the uppermost unit of the UMG), but

most commonly the Red Pine Shale lies directly beneath the Mississippian Madison

Limestone (Munroe, 2005). The uplift and erosion of the mountain range deposited

loosely cemented gravel on the mountain flanks, which overlies pre-Cenozoic bedrock.

Due to the weak cementation of sandstone and conglomerate beds on the south flank,

there have been several episodes of mass wasting in the tributaries of the glacial valleys.

Other than a few small mafic dikes (which have been dated using Rb/Sr wholerock and

K/Ar methods at 552 ± 17 to 453 ± 29 Ma) located at the crest of the uplift, igneous rocks

are almost completely absent from the Uintas (Munroe et al., 2005). Deformation has

compressed the Uintas into an asymmetrical doubly plunging anticlinal uplift (Munroe et

al., 2005).

While the Uinta Mountains do not currently contain glaciers, the topography,

dominated by deep glacial valleys, is a clear sign that these mountains were once the

home to an extensive system of alpine glaciers. Laabs and Carson (2005) conjecture that

during the Last Glacial Maximum (LGM), mountain glaciers occupied 1460 km2 in the southern part of the range. Using equilibrium line altitude analysis, Munroe and

Mickelson (2002) suggest that temperatures during the LGM were 5.5 to 8° C colder than present conditions. Deglaciation of the northern and eastern valleys of the Uinta

Mountains began around 22 to 20 ka, while deglaciation of the southern and western valleys did not begin until 18-16.5 ka (Laabs et al., 2009). Thackray (2008) adds that the glaciers did not reach their maximum extent at the western end until after the global

LGM. This difference could reflect the strong precipitation gradient (with accumulation

14 around 1000 mm per year greater at the western-most end of the range) created by the proximity of Lake Bonneville. The heightened precipitation likely caused the increased stay of glaciers in the Uintas, which was then ended by the abrupt drop in Lake

Bonnevilles water level (Munroe et al., 2006).

The current conditions in the Uintas are both warmer and drier than during

Pleistocene glaciation. Using a series of SNOTEL data collection sites, Munroe (2003) found that current mean annual precipitation ranges between 550 and 925 mm, with 60% of that coming from snow accumulation at elevations of 3300 m above sea level or greater. The mean summer temperature currently ranges from 8.7º C to 11.2º C (Munroe and Mickelson, 2002).

EJOD Lake

EJOD Lake is situated in the Uinta Mountains at the head of the West Fork of the

Blacks Fork, a tributary of the Green River (Fig. 2.1) (Munroe and Bigl, 2009). Located above modern treeline at 3323 m, it is a small lake, which reaches a maximum depth of

4.3 m and has a surface area of 2.7 ha (Fig. 2.2, 2.3). Located far up in the mountains, the surrounding landscape is dominated by a large cirque containing an apron of talus that is currently creeping downslope (Fig. 2.4). Approximately 500 m upslope from the lake is a series of end moraines that reach about 50 m higher in elevation and which, according to lichen dating, stabilized ca. 1600 BP (Munroe, 2002; Munroe and Bigl,

2009). Extending downslope from the moraine, dry meltwater channels lead into EJOD

Lake, demonstrating a direct connection between glacial runoff and lake sedimentation

(Fig. 2.5, 2.6) (Munroe and Bigl, 2009).

Figure 2.1 shows a digital elevation model of the Uinta Mountains, Utah denoting the location of EJOD Lake.

Figure 2.2. Bathymetric map of EJOD Lake (Bathymetry courtesy of T. Martel).

Figure 2.3. Topographic map of the West Fork of the Black’s Fork (contour interval =50 m, grid=1 mile; map data taken from TerraServer at http://www.terraserver.com).

Figure 2.4. Looking up into the cirque from EJOD Lake the glacial landscape is clear and most notably dominated by a glacial cirque, talus field, and a series of moraines (photo courtesy of J. Munroe). 17

Figure 2.5. A picture taken looking up from the lake towards the end moraines shows one of the meltwater channels that runs into EJOD Lake (photo courtesy of J. Munroe).

Figure 2.6. A map of the surrounding topography shows the glacial cirque, talus field and moraines. Also mapped are the meltwater channels which flow from the end moraines into EJOD Lake (photo imagery from Google Earth).

Previous studies of Glacial History in the Rocky Mountains

Previous studies of the Uinta Mountains have focused on glacial geomorphology, providing a fairly thorough understanding of glacial landforms and topography. In 1909,

W.W. Atwood published a study of the glacial geomorphology in the Uintas in his monograph Glaciation of the Wasatch and Uinta Mountains (Munroe, 2005). Atwood

(1909) found evidence of two major glacial advances, which were further studied by

Bradley in 1936. Bradley identified three periods of glacial advance, which he named the

Little Dry, the Blacks Fork and the Smiths Fork after the valleys where their sediments were best preserved (Munroe, 2005). In 1965, research by Richmond correlated these glacial advances with past climate reconstructions elsewhere in the U.S. finding that the

Little Dry glacial advance generally correlates with the Illinoian and pre-Illinoian in the

U.S. Midwest, the Blacks Fork correlates with the Illinoian, and the Smiths Fork correlates with the late Wisconsin (Munroe et al., 2005). W.R. Hansen also contributed greatly to the study of the Uinta Mountains, making observations of the formation of the regional landscape over the Middle to Late Cenozoic (e.g., Hansen, 1965, 1984, 1986)

(Munroe, 2005).

More recently, Munroe (2001) mapped the glacial geomorphology of the northern

Uintas at 1:24,000 scale. Munroe (2001) also completed a Quaternary glacial reconstruction for the area using the presence of glacial landforms to better understand the extent of valley glaciers at the time of the Last Glacial Maximum (LGM).

Subsequently, Laabs (2004) began mapping the south slope of the Uintas at a scale of

1:24,000, and by 2007 had mapped moraines, erosional trimlines, till sheets and other glacial landforms in the area. The mapping of these landforms subsequently supported

20 the reconstruction of ice extents in the Uintas during the LGM (Laabs et al., 2007).

Laabs and co-workers also used cosmogenic surface exposure dating to better understand the timing of the LGM (Munroe et al., 2005; Laabs et al., 2009).

Past studies have also investigated the more recent Holocene glacial history of the

Rocky Mountains focusing in Utah on the La Sal Mountains and the Wasatch Mountains.

Richmond (1962) studied the La Sal Mountains, using soils to identify two different ages of till and rock debris formed by Neoglaciation. He dated one unit of till as having formed between 3100 and 2800 BP and the other between 1800 and 1000 BP (Burke and

Birkeland, 1984). More recent studies have found Richmonds ages to be slightly young but the occurrence of the events is still accepted (Burke and Birkeland, 1984). In the

Wasatch Range, Ives (1950) studied moraines adjacent to cirque headwalls, which were built during the Little Ice Age. McCoy (1977) and Currey (1979) then used radiocarbon dating to provide age constraints for these tills (Burke and Birkeland, 1984).

III. METHODS

Field Work

Four overlapping cores (a surface core, two Livingstone cores, and a percussion

core) were retrieved from EJOD Lake at 40° 45.047 N, 110° 40.578 W, WGS84 on

August 12, 2005 by J. Munroe. The cores were collected from an anchored platform

(Fig. 3.1). They were then transported to Middlebury College where they were stored in

a cooler at 4°C. Samples were taken from each centimeter and placed in capped vials to

be used for bSi, C:N, and phosphorus analysis.

Figure 3.1. Cores were retrieved from EJOD Lake using an anchored platform (photo courtesy of J. Munroe).

Previous Work on EJOD Lake Core

Prior to this study, J. Munroe performed several analyses on the EJOD Lake core.

Tests were run on color, loss-on-ignition, bulk density, and grain size. The core was dated using AMS radiocarbon dating of terrestrial macrofossils to support a depth-age model. Additionally, lichen dating was performed on a series of end moraines located

500 m upslope from the lake.

Sample Analysis

Biogenic Silica Analysis (bSi)

Biogenic silica analysis was performed at 1-cm resolution to determine the amounts of organically derived SiO2 produced from diatoms in the EJOD Lake core.

Methods followed the protocol outlined by Corbett (2007). Samples were dried for 48 hours in an oven, and then ground by hand using a mortar and pestle. Between 13 and 17 mg of sediment from each sample was weighed and prepared for analysis. A 0.1M

NaOH solution was added to each sample before being placed in an 85°C shaker bath for five hours. At the end of each hour, 500 µL of sample were withdrawn and 4.5 mL of nanopure water were added. At the end of the fifth hour, molybdate reagent and a reducing solution were added to create a blue color development. The reagents were also applied to a set of standards of known silica content to create a standard curve. The colors developed for no less than 3 and no more than 5 hours before being read in a

Thermo spectrophotometer. Results were then analyzed to discriminate between biologically-derived SiO2 and mineral SiO2. A linear regression, signifying the dissolution rate of mineral SiO2, was applied based on the data points for hours 2 through

5, and the y-intercept was taken to be the total amount of biogenic silica present in the

23 sample (Fig. 3.2). Three randomly chosen replicate solutions were analyzed per run.

Additionally, any data that produced an R2 value below 0.9 were also replicated.

Figure 3.2. Excel analysis was used according to the DeMaster (1981) method to differentiate biologically derived SiO2 from mineral SiO2. The blue point represents biogenic and mineral silica, which was leached out during the first hour. The following four points represent the dissolution of only mineral silica (as all biogenic silica dissolved in the first hour). Thus a linear regression is fit to points 2-5 to represent the dissolution rate of mineral silica. The y-intercept is taken to be the total amount of biogenic silica in the sample.

Phosphorus

Phosphorus analysis was done at 2-cm resolution to determine the relative amounts of organic, occluded and mineral phosphorus in each sample. The sequential extraction, performed to geochemically distinguish the three pools of P, was adapted from the protocol outlined by Filipelli (2004). Samples were dried for 48 hours in an oven, ground by mortar and pestle and then approximately 0.1000 g of each sample was weighed and placed in a 15-mL polyethylene centrifuge tube with a plug cap. Ten mL of citrate dithionite bicarbonate (CDB) solution was added to the centrifuge tubes, which

24 were put in the shaker bath for 6 hours and then in the centrifuge for 6 minutes.

Supernatants were decanted into acid-cleaned HDPE bottles and saved for further analysis. This process was repeated adding MgCl2 and then Mill-Q water in place of

CDB, this time shaking for only 2 hours. ICP-AES was used on the supernatants at the end of this process to determine the iron-bound P concentrations. Once the iron-bound P

(occluded) had been removed, the authigenic/biogenic P was found by buffering the samples with Na-acetate and rinsing twice with MgCl2 (again shaking and centrifuging after each rinse). Ten mL of HCL was then added and the sample was shaken, centrifuged and decanted to isolate detrital P. Finally, organic P was isolated using 1 mL of MgNO3. A Thermo spectrophotometer was used to determine authigenic/biogenic, detrital, and organic P concentrations for the sequential extraction using a blue color development. Authigenic/biogenic and detrital P concentrations were then added together to determine the total mineral P concentration. For each run 4 samples were randomly chosen for replicate analysis.

C:N analysis

C:N analysis was performed at 1-cm resolution to determine the relative concentrations of carbon and nitrogen in each sample. Analysis was performed on samples which had been dried for 72 hours in an oven and then ground by mortar and pestle. To prepare for analysis, 10-20 gm of sample were weighed out on a microbalance and placed in small tin cups. Weights were recorded and later used to determine concentrations. To prevent sediment leaking, the tin cups were folded and compressed before being loaded onto a carousel and run through an elemental analyzer. To ensure accuracy, each run included a blank tin cup at the start of the run as well as a standard

(2.29% ±0.07 C, 0.21% ±0.01 N) at the beginning and end of the run and at 31 sample intervals. Additionally, 10 samples were randomly chosen for replicate analysis.

IV. RESULTS

Age Control

Four radiocarbon ages as well as -55 BP (corresponding to AD 2005 when the core was taken) for the core surface were used as age controls (Table 4.1). A depth-age model was fitted through these five points using a simple polynomial through the midpoints of the radiocarbon dates (Fig. 4.1). A thick sand layer which likely accumulated as a single rapid event was removed for the calculations, revealing a nearly linear depth-age model.

CoreDepth CorrectedDepth(cm) Material Carbon14Age CalendarAge Error (cm) (withoutsandlayer) (midpointof calibratedage) 0Ͳ55 0 38 38 pineneedle 630 607 58 94 94 pineneedle 1870 1801 89 194 181 pineneedle 3640 3963 127 234 221 pineneedle 4560 5187 136 Table 4.1. AMS ages collected by J. Munroe were used to determine the depth-age model.

Figure 4.1. A depth-age model was fitted for EJOD Lake using a simple polynomial through the midpoints of four radiocarbon ages along with -55 (corresponding to AD 2005 when the core was taken).

Detrending

The climate proxies all showed evidence of both a broader climate shift as well as superimposed smaller-scale variations. To objectively identify these variations, regressions were fit to the proxies that reflected the broader climate shifts. These regressions were then subtracted from the data, identifying deviations from the local trend (Fig. 4.2). Intervals were identified where values fell above and below the local trend, and pronounced intervals were identified as times when values fell more than one standard deviation off the local trend.

Biogenic Silica

The biogenic silica time series shows an overall decreasing trend over the past

5000 years with values reaching as high as 8.3% at approximately 4800 BP, and as low as

3.4% at approximately 1300 BP (Fig. 4.2). The trend decreases steeply between 5300 and 1400 BP (at an average rate of 0.9% every one thousand years), before flattening out almost completely between 1400 BP and AD 2005 (with only a 0.1% decrease over one thousand years). The mean bSi concentration between 5300 and 3200 BP is 6.5%.

Between 3200 and 1400 BP, the mean concentration decreases to 4.8%. The concentration continues to decrease between 1400 BP and 100 BP, with a mean of 4.6%.

Superimposed on this decreasing trend are several pronounced low intervals. To enhance these, bSi data were detrended, assuming a linear decrease in concentrations over time that suggests a broader climate shift towards less abundant diatom populations.

BSi concentrations fall below the local trend at 5330-5170, 4600-4300, 4130-4020, 3770-

3280, 2760-1620, 1380-830, 810-560, and 400-180 BP. Concentrations show

28 pronounced low intervals, falling more than 1 standard deviation below the local trend, at

5300-5200, 3570-3380, 1010-830, and 400-270 BP.

The biogenic silica time series also shows an unprecedented 8% increase in concentration between 160 BP and AD 2005. In the top-most centimeter (corresponding to AD 2005), bSi reaches a concentration of 11.5 %. Prior to 160 BP, the highest concentration reached throughout the core is 8.3%.

Figure 4.2 shows the % bSi time series for EJOD Lake. A line in the % bSi time series denotes the linear function fit to the data for detrending. A detrended data plot identifies times when % bSi values fell below the local trend.

Phosphorus

The total phosphorus measured in the EJOD Lake core shows fairly steady values with a mean of approximately 30 ȝmol g-1 throughout. Superimposed on these values are century to millennial-scale variations which become more pronounced around 2400 BP, a trend that is maintained throughout the rest of core (to AD 2005) (Fig. 4.3). The overall phosphorus time series shows no dominant trend so detrending was not used, but values do rise above the mean value at 5300-4600, 3100-2900, 2700-2100, 1600-800 BP, and

AD 1940-1970.

The percent of total P that is mineral-bound exhibits a broad rising and falling trend. Values rise between 5300 and 3300 with a mean value of 21.0%. Between 3300 and 1000 BP values attain a mean of 34.1% as they reach the peak of their rising and falling trend. During the millennium between 1000 BP and AD 2005, values initially decline and then continue to rise obtaining a mean value of 25.4%.

Superimposed on this gradual rising and falling trend are three pronounced high intervals at 3100-2700, 2400-2100, and 1700-1200 BP (Fig. 4.4). Detrending the data with a second order polynomial fit to the overall rising and falling trend reveals that % mineral P rises above the local trend at 3100-2700, 2300-1900, 1400-1000 BP, and 230

BP- AD 2005 (and rises more than 1 standard deviation above the local trend at 1890-

1560, 1380, and 180 BP).

Like the bSi data, the mineral P time series shows a dramatic shift in conditions between 160 BP and AD 2005 marked by an increase in % mineral P from 28% to 41%.

While concentrations of mineral P reach as high as 43% earlier in the data set (around

2100 BP), this is a more abrupt shift than otherwise found in the time series.

Figure 4.3. shows the total phosphorus time series for EJOD Lake.

Figure 4.4 shows the % mineral P time series for EJOD Lake. A line in the % mineral P time series denotes the second order polynomial function fit to the data for detrending. A detrended data plot identifies times when % mineral P values rise above the local trend.

C:N

The C:N time series shows an overall decreasing trend with several superimposed high points (Fig. 4.5). Between 5300 and 4200 BP, the C:N ratio has a mean of 16.1. It then undergoes a significant decrease between 4200 and 4000 BP, dropping from 16.6 to

12.7. Between 4000 BP and AD 2005 values remain low with a mean value of 12.2.

Another abrupt decrease (from 11.5 to 8.7) occurs between AD 1920 and AD 2005.

The C:N time series reaches pronounced high points at 4200, 2150, 1700, 750 and

150 BP. Assuming that the abrupt shift, which occurs between 4200 and 4000 BP, represents a significant change in the lake environment and conditions, the data between

4000 BP and AD 2005 were isolated to identify fluctuations in the record following the shift. This section of the time series was detrended using a linear regression representing a broader shift towards greater input of aquatic organic matter. Detrending reveals that the C:N ratio rises above its local trend value at 4000-3800, 2800-1300, 900-800, and

300-130 BP (reaching more than 1 standard deviation off the local trend at 1890-1560,

1380, and 180 BP).

Figure 4.5 shows the C:N time series for EJOD Lake. A line in the C:N time series denotes the linear function fit to the data between 4000 BP and AD 2005 for detrending. A detrended data plot identifies times when the C:N ratio rose above the local trend.

Sediment Flux

Using bulk density data in conjunction with a calculated annual sedimentation rate from the depth-age model, the organic, biogenic, and clastic flux were calculated in grams per cm2 per year (Fig. 4.6). The total flux shows an overall increasing trend starting at around 0.018 gm/cm2/yr at 5300 BP and reaching 0.049 gm/cm2/yr at 620 BP

(the youngest date for which flux data are available). This is an overall increase of 270%.

A linear regression, which was fit to the broadly increasing flux time series and used to detrend the data, showed no prominent departures from the local trend.

Figure 4.6 Overall flux, clastic flux, biogenic flux, and organic flux data for EJOD Lake. Solid lines for the overall flux and clastic flux time series indicate linear regressions which were fitted to the data and used for detrending.

The organic flux shows an overall decreasing trend from 0.0030 gm/cm2/yr at

5330 BP (the bottom of the core) to 0.0009 gm/cm2/yr at 3460 BP (Fig. 4.6). Between

3460 and 620 BP, organic flux shows a gradual increase from its all time low of 0.0009

gm/cm2/yr to 0.0041 gm/cm2/yr. The trends are marked by small-scale oscillations, but

are linear in nature.

The biogenic flux is slightly more complicated but can be broken into three parts

(Fig. 4.6). Between 5330 and 3520 BP, the flux is relatively constant and reaches its

lowest values of the time series with a mean value of 0.0011 gm/cm2/yr. Between 3520

and 1800 BP, the flux again remains low, although slightly higher and slightly more

constant than in the previous section. Flux in this section has a mean value of 0.0012

gm/cm2/yr. Between 1800 and 620 BP, the biogenic flux shows a gradual increase,

reaching its highest value of 0.0029 gm/yr at 1150 BP. This section has a mean biogenic

flux of 0.0018 gm/yr and is characterized by the most variable values. Detrending the

biogenic flux with a linear regression (representing a shift towards an increased

sedimentation rate), reveals century-scale variation with distinct lows at 5330, 4070, and

3520 BP and highs at 5140, 3930, and 3740 BP.

The clastic flux shows an overall linear increasing trend between 5330 and 620

BP (Fig. 4.6). Values become more variable at 2530 BP (a trend that continues to 620

BP) but continue to follow the increasing trend. Values show an overall increase of

0.033 gm/cm2/yr throughout the time series, an almost tripling of values. Distinct high

points identified by detrending using a linear regression occur at 5140, 3740, 3490, and

790 BP, and values rise more than 1 standard deviation above the local trend at 4960-

4450, 1530-1150, and 800-620 BP.

Overall Time Series

Summary statistics for bSi, mineral P, and C:N for of the EJOD Lake record

(delineated into four zones based on the three distinct zones in the mineral P record as well as the distinct shift between 182 BP and AD 2005 seen in all three proxies) are listed in Table 4.2 and shown on Fig. 4.7. Between 5300 and 5100, the bSi concentration is below its median, reaching a low point at 5260 BP. Total P shows a similar fluctuation, exhibiting a high interval between 5300 and 4600 BP. Meanwhile, between 5300 and

4600 BP mineral P gradually increases, and C:N shows century-scale fluctuation but does not reach a pronounced high or low point. BSi reaches a temporary high around 5000 BP before again falling below the local trend between 4500 and 4300 BP. The % mineral P continues to increase during this time and the C:N ratio continues to fluctuate but remains at an overall high.

Following a brief high point at 4200 BP, the bSi concentration remains low between 4100 and 4000 BP. At 4200 BP, the C:N ratio drops abruptly until 3500 and then more gradually until it reaches a local low point at 3350 BP. During this time, mineral P concentrations continue to rise. BSi concentrations reach a temporary low point between 3770 and 3280 BP, an interval which is slightly offset by a high point in mineral P between 3100 and 2700 BP. Total P is predominantly synchronous with mineral P during this interval, reaching values above its local trend between 3100 and

2900 BP.

As mineral P concentrations continue to drop, bSi concentrations reach a temporary high at around 2900 BP before falling below the local trend between 2560 and

1620 BP. Beginning shortly after the onset of the bSi low interval, C:N reaches a temporary high at 2000 BP which lasts until 1700 BP. Starting at 2300 and lasting until

1900 BP mineral P concentrations reach a temporary high point, which is synchronous with a high interval in total P. Mineral P concentrations then begin to decrease between

1400 and 1000 BP while bSi concentrations decrease between 1380 and 1240 BP.

During this time, C:N continues to oscillate, decreasing slightly from a temporary high reached at 1700 BP.

At 810 BP, bSi concentrations return to below mean values, where they remain until 560 BP. During this time, C:N reaches a temporary high at 750 BP, but does not surpass the local trend, and mineral P concentrations continue to decrease. At 640 BP,

C:N rises above its local trend where it remains until 310 BP. Immediately following this temporary high period, bSi concentrations fall below their mean value where they remain until 180 BP, and mineral P concentrations rise above their mean value at 230 BP and remain above until AD 2005. At 180 BP, all proxies show a dramatic shift, characterized by steeply increasing bSi and mineral P values and decreasing C:N ratios (Fig. 4.7).

Zone Age Mean StandardDeviation CoefficientofVariation Zone4 180BPͲAD2005 bSi6.5% 2.7 41.1% MineralP 29.5% 11.3 38.1% C:N11.5 1.0 8.6% Zone3 1000Ͳ180BP bSi4.5% 1.1 24% MineralP 23.5% 6.0 25.3% C:N11.7 0.7 6.0% Zone2 3300Ͳ1000BP bSi4.8% 1.1 23% MineralP 34.1% 11.5 34.9% C:N12.4 0.7 5.5% Zone1 5300Ͳ3300BP bSi6.4% 1.5 23.0% MineralP 21.0% 6.6 31.0% C:N14.6 1.8 12.6% EntireCore 5300BPͲAD2005 bSi5.4% 1.6 30.0% MineralP 27.8% 11.0 39.5% C:N12.9 1.7 13.0% Table 4.2. Summary statistics for 4 subzones in the EJOD Lake core.

Figure 4.7 Summary of C:N, % mineral P, and % bSi data.

V. DISCUSSION

Relationships between Proxies

Recently analyzed proxies were compared with previously collected analysis on loss-on-ignition, bulk density, median grain size, and color (Fig. 5.1). An a* + b* value was calculated to represent increasing pinkness in the sediment (where high a* values suggest more red color in sediments and high b* values suggest more yellow color in sediments).

Figure 5.1. shows a comparison of proxies analyzed in this study with bulk density, % LOI, median grain size, and a*+b* data previously collected by J. Munroe.

BSi and LOI

Both biogenic silica and loss-on-ignition are measurements of organic matter recorded in the sediment core. However, while bSi records the abundance of aquatic organic matter derived from diatoms, the LOI values are a measurement of both aquatically and terrestrially derived organic matter. It is thus useful to compare these two proxies, to get a better understanding of how overall organic sedimentation varies with diatom populations in the lake. Furthermore, the inwashing of terrestrial aquatic matter has the ability to impact the diatom populations in the lake, and thus it is interesting to inspect whether increased LOI values correspond with increases or decreases in biogenic silica abundance.

Throughout the core, bSi and LOI are closely aligned as they follow a gradual, linearly decreasing trend (Fig. 5.1). It should be noted however that the LOI trend decreases more rapidly with a slope of 0.0015 % per yr while bSi decreases with a slope of only 0.0007 % per yr, less than half the slope of LOI. The general decreasing trend however in LOI can be divided into two sections: 5330 to 3990 BP, and 3990 BP to AD

2005. In the older section (5330 to 3990 BP), LOI shows a steeply decreasing trend with a linear slope of 0.0018 % per year, while in the younger section (3990 BP to AD 2005)

LOI follows the bSi trend much more closely with a slope of 0.0008 % per year.

Additionally, the two data sets show aligned oscillations with significant lows seen in both data sets at approximately 5200, 3450, 1900, and 620 BP (Fig. 5.1).

This alignment between the two proxies suggests that bSi and LOI are strongly correlated. Between 3990 BP and AD 2005, where the two proxies show very similar decreasing trends as well as smaller-scale oscillations, it appears that most of the organic

44 matter being recorded in the lake is of aquatic origin. Before 3990 BP however, where bSi and LOI are positively correlated, higher LOI values indicate that there is a greater amount of terrestrial organic matter being washed into the lake.

BSi and C:N

The comparison of C:N data to bSi concentrations is interesting as both proxies can be used to indicate the relative abundance of aquatic organic material. C:N ratios in isolation are somewhat difficult to interpret as it can be hard to tell whether rises and falls are driven by changes in C, N or both. However, in conjunction with bSi, which indicates the relative abundance of aquatic organic material, C:N ratios can be quite revealing. Given that higher C:N values indicate a greater abundance of terrestrial organic material, an inverse relationship between C:N and bSi would be expected.

In the EJOD Lake core, bSi and C:N ratios are in fact inversely related (Fig. 5.1).

This is seen in the low intervals in biogenic silica at 3770-3280, 2760-1620, and 810-560

BP which are accompanied by high intervals in C:N. Pronounced lows in bSi (values more than 1 standard deviation off the local trend) generally precede high intervals in

C:N by one to two centuries, while high intervals in C:N continue two to three centuries after bSi has returned to its local trend (Fig. 5.1). This inverse relationship suggests that decreased lake productivity leads to periods of lake sedimentation that are more dominated by terrestrially-derived organic matter.

LOI and C:N

Used in conjunction, LOI and C:N data can help determine whether the organic material measured in LOI analysis is of terrestrial or aquatic origin. The dramatic shift in

C:N between 4200 and 4000 BP occurs at virtually the same time as the decreasing LOI

45 trend becomes less steep (Fig. 5.1). Between 5300 and 4200 BP, high C:N values as well as high LOI concentrations suggest that terrestrial organic matter dominated organic lake sedimentation. After 4000 BP however, lower C:N and LOI values indicate that aquatic organic matters plays a much more dominant role. The stronger correlation between bSi and LOI concentrations after 4000 BP further supports this theory.

BSi and Median Grain Size

Diatom populations are strongly influenced by both water temperature and the amount of sediment that is being washed into the lake. Because diatoms are photosynthetic, increased lake cloudiness due to the inwashing of sediments can lead to decreased diatom populations, measured as drops in bSi concentrations. It is therefore interesting to investigate the effects of lake sediments (represented by median grain size measurements) on bSi concentrations.

Throughout the EJOD Lake record there is a negative correlation between % bSi and median grain size, where intervals of less abundant diatom populations correlate with above mean value median grain sizes (Fig. 5.1). This relationship holds true throughout the entire time series except for the interval between 340 BP and AD 2005 where both % bSi and median grain size are above their local trends. It should be noted however that while they are both above their means, biogenic silica continues to decrease as median grain size increases.

This inverse relationship implies that increases in median grain size correspond to decreases in biogenic silica. This could be because larger grains being washed into the lake reflect increased clastic input which could cause greater cloudiness or turbidity in the lake and make it harder for diatom populations to flourish.

C:N and Median Grain Size

The relationship between C:N and median grain size can help indicate whether grain size is larger when sedimentation is dominated by terrestrial organic material or when it is dominated by aquatic organic material. In EJOD Lake, at 5170, 1960 and 200

BP, peaks in median grain size correlate with peaks in C:N suggesting that grain size is larger when organics are being washed off the landscape into the lake (Fig. 5.1). For the rest of the record however, there is a negative correlation, suggesting that often grain size decreases when the organic matter is dominantly terrestrially-derived.

Mineral P and Median Grain Size

A obvious meltwater channel running from upslope moraines into EJOD Lake make mineral P a good proxy for periods of increased inwash of glacially eroded sediments. Consequently, a comparison between % mineral P trends and median grain size can help reveal the nature of the eroded sediments.

The EJOD Lake record shows four periods where mineral P is elevated above its local trend, implying heightened periods of inwash of phosphorus that was eroded directly from the bedrock. During the three earliest periods, from 3100 to 2700, 2300 to

1900, and 1400 to 1000 BP, there appears to be a correlation between elevated percentages of mineral P and lower median grain sizes (Fig. 5.1). In fact, during these three periods, median grain size is predominantly below its mean value. This correlation suggests that sediments eroded off the bedrock by glaciers and eventually washed into the lake were dominated by very fine-grained rock flour. During the most recent interval of elevated mineral P however, from 230 BP to AD 2005, median grain size remains above its mean for the duration. This interval of increased mineral P however likely reflects

47 more recent anthropogenic affects on the landscape as no glacier was present in AD 2005 to erode the bedrock, thus explaining the lack of correlation.

Mineral P and Biogenic Silica

Both % bSi and % mineral P can be indicative of periods of Neoglaciation. While decreased temperatures and increased erosion into the lake during periods of

Neoglaciation create declines in % bSi, they also create increases in % mineral P.

Further, the increased erosion indicated by intervals of high mineral P leads to less ideal conditions for diatom populations. Consequently, it would be expected that decreases in

% bSi are seen in conjunction with increases in % mineral P.

As expected, there is a strong correlation between % mineral P and % bSi. In fact, the four notable highs in % mineral P all follow a low in bSi. The most pronounced lows in % bSi, as indicated by values more than 1 standard deviation below the local trend value, are all followed within one to three centuries by a pronounced high in % mineral P (Fig. 5.1). This offset likely indicates that much of the mineral P is being washed into the lake during glacial retreat. It follows that during periods of

Neoglaciation, cool temperatures cause decreases in diatom populations, and then as temperatures warm and the glacier retreats, diatom populations again flourish as fine rock flour enriched in mineral P is washed from the moraines into the lake. The low intervals in biogenic silica occurring between 5330 and 5170, 4600 and 4300, and 4130 and 4020

BP, however, do not correspond to an increase in % mineral P above its local trend.

However, while it does not rise above mean values, % mineral P does show local maximums at 5200, 4250, and 3690 BP. These three maximums all occur within a few hundred years of a local minimum in % bSi, showing a similar offset relationship. The

48 interval between 890 and 560 BP, marked by below mean % bSi values, is the only low period which does not correspond to a high interval in % mineral P. Additionally, the offset between intervals of low biogenic silica and high mineral P indicates that temperature, not turbidity, is the main control of biogenic silica in EJOD Lake.

Mineral P and C:N

C:N and mineral P are both indicators of terrestrial materials being washed into the lake. While higher C:N ratios indicate greater inwashing of organic matter, higher % mineral P indicates increased amounts of bedrock being eroded and washed into the lake.

As previously indicated, high intervals in C:N correspond to low intervals in % bSi.

Therefore, it would be expected that high intervals in C:N also correspond with high intervals in % mineral P. Not surprisingly, there is in fact such a correlation (Fig. 5.1).

Further, the abrupt drop in C:N which occurs between 4200 and 4000 BP occurs during a low interval in biogenic silica which lasts from 4130 to 4020 BP, and shortly before a local high in % mineral P at 3690 BP. This suggests that this abrupt drop occurred during a time of Neoglaciation.

Mineral P and Total P

While the total P is a measure of the amount of phosphorus that is in the lake system, the mineral P indicates what percent of that total is mineral-bound. While the total P shows an overall decreasing trend between 5300 and 3500 BP, % mineral P increases during this interval. This suggests that while the total P in the system was decreasing, its overall composition was becoming more and more dominated by mineral

P. Mineral P and total P however do show some synchronicity between 3100 and 800

BP, as the three distinct high intervals in % mineral P correspond with high intervals in

49 total P as well (especially the two more recent intervals at 2400-2100 and 1700-2100

BP).

Color, Clastic flux, and Mineral P

While color analysis on its own is quite hard to interpret, it is a good indicator of visible changes in the sediment record. For EJOD Lake, a*+b* data were calculated to represent times of increased pinkness in the sediment color. Given that the bedrock around EJOD Lake is dominated by pink quartzite, color analysis is likely a good indicator of the true color of stained quartz grains being washed into the lake.

The four dominant high intervals in clastic flux (identified by detrending) at 4960-

4450, 1530-1150, and 800-620 BP all align with periods of increasing a*+b* values (Fig.

5.2). A*+b* values also align well with peaks in mineral P concentrations, reaching local high points during the three major high intervals of % mineral P at 3020-2860, 2300-

2050, and 1400-1210 BP. This indicates that sediment is pinker during times of increased glacial erosion and inwashing of clastic material into the lake.

Figure 5.2 shows a comparison of a*+b* data (representing the relative pinkness of the sediment) with % mineral P and clastic flux.

Interpretation of EJOD Lake History

The EJOD Lake record shows strong evidence of a gradual cooling trend. The oldest data present, from 5300 BP, indicate that during the middle Holocene, EJOD Lake area was characterized by a mild climate and abundant vegetation. High values in C:N between 5300 and 4200 BP indicate that lake sediment during this early interval was dominated by terrestrial organic matter. This, coupled with the fact that % LOI is much

51 greater than % bSi, implies that a large amount of organic matter was being washed off the landscape and into the lake. Decreasing bSi and LOI concentrations as well as increasing bulk density values throughout the record, however, show a gradual shift into

Neoglacial conditions with a climate less suitable for abundant diatom populations and organic matter growth. This theory is supported by upward trending clastic flux and overall flux values which show a transition to a more clastic dominated system. Cooler climates would have produced a less hospitable environment for organic matter as well as created the more intense freeze-thaw weathering climate indicated by such increases in clastic flux.

Superimposed on this long-term cooling trend are several century-scale variations which represent periods of Neoglacial advance and retreat (Fig. 5.3, 5.4). The interval between 5300 and 5200 BP is marked by very low organic activity, indicated by a dip in both % bSi and % LOI. The sediment input during this time was dominated by fine- grained clastic input indicated by a high interval in clastic flux and a dip in median grain size, followed by a slight increase in mineral P. This clastic dominated environment reflects an early Neoglacial advance, where harsh temperatures limited diatom populations and phosphorus-rich rock flour created by glacial erosion was washed into the lake. A slight increase in C:N at this time may indicate a greater amount of organic matter being washed into the lake by glacial erosion as well.

Figure 5.3 A multi-proxy study of EJOD Lake shows evidence of several periods of Neoglacial advance.

Figure 5.4. A summary table of notable intervals for the EJOD Lake record. Colored boxes indicate significant intervals (as indicated in the right column), which suggest Neoglacial activity. Darker colors indicate times when data fell more than 1 standard deviation off the local trend.

At around 5200 BP, % bSi values begin to increase, eventually reaching a peak at

4700 BP (% bSi values do not get this high again until around AD 1950). These increasing bSi values show the reemergence of a dominant diatom population as conditions return to the warm temperatures which marked the early to middle Holocene.

This shift is accompanied by increasing % LOI values, which remained significantly higher than bSi concentrations during this time, and continued high C:N ratios which indicate that organic sediment remained dominantly terrestrially derived. This likely reflects a return to the abundant vegetation conditions which characterized the early and middle Holocene. Throughout this interval however, mineral P concentrations continued to increase showing the overall trend of a shift towards a harsher, more intense weathering climate.

These trends continue until around 4600 BP when % bSi began to decline, falling below its mean value at 4600 BP. This drop is accompanied by a decrease in LOI concentrations. BSi and LOI remained synchronous showing a strong correlation between lake productivity and the inwash of terrestrial organics. C:N reached its highest values for the entirety of the core at 4200 BP, likely the result of both a decrease in in- lake productivity, and high amounts of terrestrial organic material being washed into the lake by weathering processes. This high in C:N directly corresponds to a peak in % mineral P as well as a rise in clastic flux. While proxies indicate a shift towards cooler conditions, bSi, C:N and mineral P did not reach pronounced lows suggesting that this interval is indicative of the larger-scale shift into Neoglacial conditions. This shift likely caused a more intense weathering climate dominated by freeze-thaw erosion processes and declining diatom and organic populations.

After reaching a high point at 4200 BP, C:N ratios dropped dramatically until

4000 BP and then continued to drop more gradually until they reach a low point at 3250

BP. This dramatic shift is much more pronounced than other oscillations in the C:N time series, and following the decline, values never return to more than three-quarters of the pre-4200 BP values. One possible explanation for this decline is that it represents a shift in treeline from a position above EJOD Lake to below the lake, which likely occurred as the result of the overall declining temperatures that marked the middle to late Holocene.

Such a shift would have greatly reduced the amount of terrestrial organic matter being washed into the lake. This interpretation is consistent with LOI and bSi relationships which established a similar, constant slope around 4000 BP. Prior to the shift, LOI values were much higher than bSi values, suggesting a large terrestrial component to the

55 organic deposition in the lake. However, at 4000 BP, LOI values flattened out and began to more closely mirror bSi values suggesting an overall decrease in terrestrial organic matter which accompanied the fall in treeline.

Following this shift, the lake temperatures began to warm and bSi and LOI values began to increase. Warmer temperatures led to the inwashing of mineral P eroded from the landscape, and a consequent increase in clastic flux.

The interval between 3600 and 3400 BP reflects a renewed period of

Neoglaciation in which colder temperatures and increased clastic input led to decreased diatom populations and low LOI values. It is in this interval, at 3500 BP, that a 15-cm flood layer occurred. It was removed because it was interpreted to represent a single depositional event. While it is possible that departures from the local trend in proxies around this point may represent artifacts from the flood layer, values show a significant departure from the local trend even with a centimeter on either side of the flood layer removed in both the bSi and LOI proxies, suggesting that this is indeed indicative of a

Neoglacial advance. The increase in clastic input is delayed and does not occur until

3500 BP, suggesting that colder temperatures lead to the decreased organic populations even before glacial inwash increases. A high interval in C:N does occur but not until

3000 BP, suggesting that temperatures cooled, leading to diminishing diatom and organic populations for several centuries before glacial erosion increased leading to a greater inwash of terrestrially-derived organic matter.

At around 3200 BP, a return of bSi and LOI concentrations to local trend values indicates an increase in local temperatures. This warm period likely also led to glacial retreat during which fine-grained, phosphorus-rich sediments were washed into the lake.

This period of increased inwash is marked by high % mineral P, low median grain size, and high magnetic susceptibility which was sustained until 2600 BP. This offset between low diatom populations and eroded P being washed into the lake indicates that the greatest inwashing of glacially eroded sediment occurred during glacial retreat.

BSi and LOI concentrations greater than 1 standard deviation below their local trends indicate another period of Neoglacial advance between 2500 and 2300 BP. This period is closely followed by a pronounced high period in mineral P concentration between 2300 and 2000 which indicates glacial retreat. During this time, bSi and LOI values returned to their local trends, as glacial meltwater causes increased inwash of P- rich glacial flour which had been stored in the moraines.

Around 2000 BP, low bSi and LOI values and high C:N values indicate that the

EJOD Lake area entered into another period of Neoglaciation. Values remained low until

1600 BP, when lichenometry indicates that the upslope moraines stabilized (Munroe and

Bigl, 2009). Ice retreat began at 1600 BP following the stabilization of the moraines, and is seen in mineral P values, which began to increase at 1650 BP and fall more than 1 standard deviation above their local trend at 1400 BP. These values are accompanied by low median grain size and high clastic flux which are further evidence of a period of glacial retreat.

Following the 1600 BP retreat, brief low intervals in % bSi and % LOI at 1500 and 1000 BP show evidence of climate variability. LOI and bSi values did not however reach pronounced lows (more than 1 standard deviation below their local trend) suggesting that this variability does not mark a full-on Neoglacial advance. Continued

57 high intervals in % mineral P, magnetic susceptibility, and low median grain size, indicating continued glacial retreat, support this theory.

Lake conditions were quite variable between 1000 BP and AD 2005. Low bSi and LOI concentrations, however, at 800-560 and 400-180 BP show signs of two cold periods during these intervals. The 800-560 BP interval is accompanied by a high in clastic flux, and the 400-180 BP interval is synchronous with a peak in C:N but neither mineral P nor median grain size show a prominent shift during these episodes.

Additionally, the youngest dated moraines upslope of EJOD Lake stabilized around 1600

BP, making it unlikely that this cold interval represents a period of glacial advance and retreat. Instead, low bSi and LOI concentrations represent periglacial activity.

Beginning at 180 BP, lake conditions show a dramatic shift, marked by an abrupt rise in % bSi, % LOI, and % mineral P, and a steep decrease in C:N ratios. BSi reached unprecedented high values, and mineral P and LOI values show their most distinct upwards shift. The C:N shift in this section resembles the dramatic drop it underwent between 4200 and 4000 BP. This sudden change in conditions in the last 250 years suggests an anthropogenic-driven shift. The increase in % mineral P certainly does not reflect increased glacial erosion as no glacier is currently present in the EJOD area. The increase in mineral P is more likely reflective of a decrease in soil formation and ground cover, exposing bare soil, which contains mineral P. A simultaneous increase in total P likely fertilized EJOD Lake, thus leading to dramatic increases in the diatom populations.

The decline in C:N values may represent both a shift towards more productive in-lake conditions leading to increased nitrogen values, as well as overuse of the land leading to the reduction of terrestrial organic matter surrounding the lake.

Links to Regional and Global Climate

Onset of Neoglaciation

The overall cooling trend indicated in the EJOD Lake record corresponds to a general decline in temperatures throughout North America marked by the onset of

Neoglaciation during the middle to late Holocene. This is consistent with a study by

Reinemann et al. (2009) of a small sub-alpine lake in Nevada which suggests that North

American climate during the Holocene consisted of a warm middle Holocene followed by cooler Neoglacial conditions with peak temperatures occurring at about 5400 BP.

The onset of Neoglaciation occurred around 5000 BP in other parts of the world as well. A study done in west-central Jotunheimen in Scandinavia found that the onset of

Neoglaciation occurred shortly before 5730 BP (Matthews et al., 2005). Glacier reconstructions from northern Folgefonna in western Norway indicate that an early phase of glacier growth occurred around 5200 BP (Bakke et al., 2005). Lake studies from northern Sweden, found the onset to be somewhat earlier at 7000 BP, but did show a gradual cooling trend, with temperature decreases of approximately 1.2qC since the early

Holocene (Bigler et al. 2002). Snowball et al. (1999) suggest that the onset of

Neoglaciation in Sweden occurred between 6100 and 5700 BP. These dates correlate to cold events in Greenland ice cores, and lake sediment records from southern Germany

(Snowball et al., 1999). An even earlier onset of Neoglacial advance has been suggested for central Tibet and the bordering mountains where glacial advance occurred between

9400 and 8800 BP (Yi et al., 2008). Holocene Neoglacial activity has also been recognized in South America where a middle Holocene climatic transition has been found

(Rodbell et al., 2008). In New Zealand, this Neoglacial advance occurred between 5400 and 4900 BP.

While the onset of Neoglaciation is not apparent from the EJOD Lake core as the successfully retrieved core only dates back to 5300 BP, the general decreasing trend seen in % bSi, % LOI, and the increasing trend in clastic flux show clear signs of a shift into

Neoglacial conditions consistent with studies in the Rocky Mountains as well as throughout the Northern and Southern Hemispheres.

These general decreasing trends are also synchronous with declining values in mean northern hemisphere June insolation (Fig. 5.5). Values follow a broadly decreasing trend with a relatively steep slope around 5000 BP (around the onset of Neoglaciation).

Insolation continues to decrease steadily until present, flattening only slightly in the last

1000 years (Berger, 1992).

Figure 5.5. A comparison of mean summer insolation data with clastic flux (on an inverted scale), bulk density, LOI, and bSi data shows strong evidence of gradually decreasing temperatures between 5300 BP and AD 2005, which correspond to decreasing insolation values (Berger, 1992).

The 4200 Drought

Between 4300 and 4100 BP, The mid-continent of North America was affected by a severe drought which resulted in rapid climate changes (Booth et al., 2005). The widespread effects of this event have been recorded throughout the central and western

United States including in the Rocky Mountains (Stone and Fritz, 2006). The EJOD

Lake time series shows evidence of high intervals in C:N, % mineral P, and clastic flux, and low intervals in median grain size, % LOI, and % bSi between 4300 and 4100 BP.

C:N also shows a steep downwards shift between 4200 and 4000 BP. While these proxies do not show values more than 1 standard deviation off their local trend, the overlap of low median grain size, % LOI, % bSi, and high C:N, % mineral P, and clastic flux, suggest a climate change that may be related to the synchronous 4200 drought. It is possible that the high intervals in clastic flux and % mineral P were produced by a change in normal precipitation that diverted a stream or other flow into the lake. This subsequent increase of inwashing sediments may have made the water cloudy, leading to a decline in diatom populations. Further, the drought may have amplified the affects of already cooling temperatures, prompting a shift in treeline to below the lake which is suggested by the dramatic drop in C:N at 4200 BP.

Periods of Neoglacial Advance

The EJOD Lake record shows a general trend of decreasing temperatures culminating in three episodes of Neoglacial advance. These episodes occurred at 3600-

3400, 2500-2300, and 2000-1600 BP (Fig. 5.4).

While Holocene Neoglacial advance is difficult to compare regionally as it is very dependent on altitude and local environment, there is evidence of similar advances

62 elsewhere in the western Rockies and globally (Fig. 5.6). For example, evidence of the

3600 to 3400 BP advance has also been found in the Elk Mountains of Colorado

(Birkeland, 1973), and the Front Range in Colorado (Benedict, 1993). These advances may be part of the Temple Lake Stade which occurred sometime between 4000 and 2000

BP (Janke, 2005). Additionally, Nicholas and Butler (1996) found evidence of periglacial rock glacier activity during this period in the La Sal Mountains of Utah. A study by Refsnider and Brugger (2007) indicates a period of Neoglacial advance associated with an interval of climatic deterioration a few centuries later at 3080 BP throughout the southern Rocky Mountains. During this interval, however, restored diatom populations and LOI values as well as high % mineral P and clastic flux indicate glacial retreat in the EJOD Lake area. There is little evidence of a synchronous episode in other parts of the northern or southern hemisphere. In a study from the southern basin of Lake Titicaca in Bolivia and Peru, the 3600 to 3400 BP interval is characterized by high organic matter and coarser grains which are indicative of a water-level rise after a prolonged mid-Holocene dry event (Abbott et al., 1997). While the 4200 drought was seen throughout much of North America, proxies in the EJOD Lake core show a decrease in organic sediments and an interval of low median grain size at this time. Thus this return to normal wetness conditions does not appear to be synchronous between EJOD

Lake and areas in the Southern Hemisphere.

Figure 5.6. A regional comparison of notable intervals for the EJOD Lake record with evidence of Neoglacial advances elsewhere in the western U.S. Darker shading in EJOD proxies indicates intervals where values were more than 1 standard deviation off the local trend.

The 2500-2300 BP cooling event suggested by the EJOD Lake time series correlates well with periods of Neoglaciation in the Uinta Mountains, Utah (Munroe,

2002), the La Sal Mountains, Utah (Nicholas and Butler, 1996), the Elk Mountains and

Sawatch Range, Colorado (Birkeland, 1973; Refsnider and Brugger, 2007), the Front

Range, Colorado (Benedict, 1993), and Gannett Peak and Black Joe, Wyoming (Dahms,

2002), which all generally correspond to glacial and periglacial activity during the

Audubon Stade (Refsnider and Brugger, 2007). A study farther north, in the Canadian

Rockies, indicates a period of glacial advance called the Peyto Advance, which occurred around 2400 BP (Pederson, 2000). Bakke et al. (2005) used physical sediment variability in glacier-fed lakes in Norway to indicate that gradual glacial expansion culminated in a glacial event at 2300 BP. However there is little other evidence of a synchronous glacial event outside of western North America. Yi et al. (2008) indicate that throughout Tibet and the surrounding mountains, there is evidence of Neoglaciation between 3500 and

1400 BP. While this does correspond with the 2500-2300 BP event identified in the

EJOD Lake core, it is a very broad range and thus hard to correlate with other studies.

The episode from 2000-1600 BP, characterized by low intervals in LOI and bSi and a high interval in C:N appears to be less regionally synchronous than previous episodes. There is some evidence of Neoglacial advance elsewhere in the Uinta

Mountains (Munroe, 2002) and the La Sal Mountains (Nicholas and Butler, 1996), but little evidence outside of Utah. There is also some evidence based on rock glacier reconstructions of a period of glacial activity (also corresponding with the Audubon

Stade) in the Elk Mountains and Sawatch Range of central Colorado between 2070 and

1150 BP. However this glacial activity is not apparent in many other climate proxies for the area (Refsnider and Brugger, 2007). There is very little evidence of synchronous glacial activity in other parts of the world.

The Little Ice Age

In many parts of the Northern Hemisphere, Holocene Neoglaciation culminated in the Little Ice Age, a period of glacial advance, which occurred between AD 1250 and AD

1850. Most Northern Hemisphere glaciers reached their maximum extent at this time,

65 destroying geomorphic evidence of previous glacial advances (Kaplan et al., 2002). In the Wasatch Range, clear evidence of moraines formed during the Little Ice Age has been studied using radiocarbon dating (Burke and Birkeland, 1984). There is also evidence of the Little Ice Age throughout the world, including data from extensive studies in northern

Fennoscandia (Korhola et al., 2000), and Tibet (Yi et al., 2008).

However there is no definite evidence of correlative glacial landforms in the northern Uintas (Munroe, 2005). Rephotography studies in the area indicate summer temperatures in AD 1870 of approximately 1qC below modern values, which would have been more favorable for the formation of small alpine glaciers (Munroe, 2003).

However, Munroe (2002) suggests that this lack of glacial activity is a result of conditions that were too dry for the formation of cirque glaciers.

The EJOD Lake time series shows evidence of climatic deterioration during the

Little Ice Age at 810-560 and 400-180 BP, periods that were characterized by low intervals in LOI and bSi concentrations. The most recent interval also shows a correlative peak in C:N, while the earlier interval is synchronous with an episode of increased clastic sediment from the surrounding watershed. However, while most North

American glaciers reached their maximum extent during the Little Ice Age, the corresponding EJOD proxy fluctuations are not as prominent as they are elsewhere in the record suggesting that the effects of the Little Ice Age were less pronounced. Further, a prominent Little Ice Age advance in the EJOD Lake area would likely have wiped out the end moraine which stabilized around 1600 BP. It could be that like many areas in the northern Uintas, a dry climate prevented the formation of a large cirque glacier at this time.

Anthropogenic Changes

Dramatic changes in several proxies between 180 BP and AD 2005, including an abrupt rise in LOI, bSi, and mineral P as well as a steep fall in C:N, show evidence of a definite environmental shift during the last 230 years which has been observed broadly throughout the world and may be related to anthropogenic activity.

While these interpretations are more speculative (due to the sampling interval, this period represents only 14 data points, and only 7 in the case of phosphorus analysis,), this dramatic shift likely represents the onset of grazing ca. AD 1850 (which given the uncertainty in the age model, corresponds reasonably well). As the state of Utah is very dependent on the production of livestock for economic means, grazing has been an interesting issue since its inception ca. AD 1850. Utah is particularly dependent on the wise-use of land as its short growing season, dry climate, and mountainous landscape make much of its land inhospitable to cultivation (Pickford, 1932). The amount of livestock in Utah grew dramatically between 1880 and 1910, and by 1931, a reported

344,000 cattle and 2,926,000 sheep had been introduced to the state. In 1880, livestock expanded from the foothills, where they were initially concentrated, into the higher mountain ranges and semi-desert areas (Pickford, 1932). Studies show that these practices have led to a definite decrease in plant cover, reducing the cover of perennial grasses by as much as 85 % (Pickford, 1932), a change which could have a strong impact on the landscape as bare land has been shown to lose soil up to 123 times faster than soil with as little as a sod cover (Pimentel et al., 2005).

The EJOD Lake site is an area of very intense grazing, and is dominated by up to

40 % bare soil in places. The introduction of sheep in the area would have exposed bare soil and increased soil erosion. This influx of bare, poorly weathered sediment which likely represents soil parent material washing into the lake explains the increasing mineral P values. During this time, the total P in the lake sediment record also increases, a characteristic which likely reflects more rapid phosphorus cycling due to the natural cycles produced by the grazing sheep. The very dramatic increases in bSi likely reflect a response to the fertilization of the lake by the inwashing of phosphorus. This overall increase in in-lake productivity, leading to more nitrogen-enriched organic sediments would similarly explain the drop in C:N ratios.

VI. CONCLUSIONS

The EJOD Lake time series shows strong evidence of a gradual cooling trend between 5300 BP and AD 2005 which corresponds to decreasing Northern summer solar insolation values in the last 5000 years and follows the broader shift throughout the

Northern Hemisphere from a warm, moderate climate which characterized the early to middle Holocene to a cooler, harsher climate. This shift is characterized by a decrease in organic input, a decline in diatom populations, and increases in total and clastic flux.

Declining biogenic silica and organic matter concentrations as well as increasing mineral P concentrations between 5300 and 3000 BP indicate a gradual shift into

Neoglacial conditions, including a dramatic shift at 4200 which may indicate a drop in tree line to its present-day position below the lake. This gradual decrease in temperatures may have been amplified by the 4200 Drought.

Superimposed on this long-term trend, EJOD Lake shows century-scale patterns of declining diatom populations and organic material as well as a shift towards more terrestrially derived organic matter followed by periods of increased mineral P concentrations and clastic flux, and low median grain sizes at several points throughout the record. These are interpreted as three periods of Neoglacial advance at 3600-3400,

2500-2300, and 2000-1600 BP. The most recent advance at 2000-1600 BP likely deposited the prominent upslope moraines which stabilized around 1600 BP. Subsequent increases in concentrations of mineral P suggest that an influx of glacial flour stored in

69 the moraines was washed into the lake during glacier retreat at 3000-2900, 2300-2000, and 1400-1200 BP.

The three identified episodes of Neoglacial advance correspond broadly to glacial activity elsewhere in western North America including advances identified in Utah,

Colorado, Wyoming, California and the Canadian Rockies. There is some synchronicity between the EJOD Lake record and identified Neoglacial advances in Scandinavia and

Tibet, but there does appear to be a strong regional dependence.

Smaller fluctuations in bSi, LOI, C:N, mineral P, clastic flux and median grain size show evidence of periglacial activity consistent with the Little Ice Age at 810-560 and 400-180 BP. While these fluctuations are not as prominent as earlier Neoglacial advances, they may still represent cold intervals and corresponding landscape change.

EJOD Lake also shows signs of a dramatic shift in conditions during the last 200 years, which is relatively synchronous with the onset of grazing in this region. New land practices likely led to a decline in vegetative cover and increased erosion, which led to the fertilization of the lake and bolstered in-lake productivity.

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