Late Holocene Glacial Advances in the Klamath Mountains

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CALIFORNIA STATE UNIVERSITY, NORTHRIDGE

LATE HOLOCENE GLACIAL ADVANCES IN THE KLAMATH MOUNTAINS,

NORTHERN CALIFORNIA, DETERMINED FROM 10BE COSMOGENIC

EXPOSURE DATING AND DENDROCHRONOLOGY

A thesis submitted in partial fulfillment of the requirements

For the degree of Master of Science

in Geology

Joshua T. Graham

December, 2013

The thesis of Joshua T. Graham is approved:

______

Dr. Julie Laity Date

______

Dr. James Hayes Date

______

Dr. John Yule Date

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Dr. Richard Heermance, Chair Date

California State University, Northridge

Acknowledgements

The completion of my thesis would not have been possible without the assistance and support of many people along the way. First and foremost, I thank my advisor, Dr.

Richard Heermance, for his extraordinary guidance, knowledge and fortitude throughout the coarse of my studies at CSUN. His passion for geology and enthusiasm about research has been an endless source of motivation for me over the last two and a half years. I also want to extend great appreciation to my committee members, Dr. Julie Laity,

Dr. James Hayes and Dr. Doug Yule for their tireless efforts to bring my thesis to the highest level of scientific integrity possible. I am very appreciative of the faculty, staff and students in the Department of Geological Sciences at CSUN for the invaluable support and encouragement.

I would also like to acknowledge my family and friends for their support during my graduate studies. Particularly, my friends and fellow classmates Ryan Witkosky and

Jose Cardona, who spent two weeks camping in the Trinity Alps to help me complete my field work in the summer of 2012. I also thank Marylyn Hanna and the Geological

Society of America for their financial contributions, which made my field and lab work possible.

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Table of Contents Signature Page……………………………..………...………..…………………………..ii Acknowledgements…………………..………..……..…………………………………..iii List of Figures…….………………..…………..……..…………………………………..vi List of Tables…….………………………..……..………………..…………………….viii Abstract……………………………..……………………………..………………….…..ix Chapter 1: Introduction……………………………………………..…………………..…1 Chapter 2: Study Area……………….……………………………..……………………...4 Chapter 3: Background……………………………………………..…………….…….…5 3.1 Glacier Response to Climate………..…………………….…………...... ……5 3.2 Holocene Climate Variability…..……………………..……..………………..6 3.3 Previous Studies……………..………………………………………….....…..9 3.3.1 LIA in the Northern Hemisphere……...…………...... …...9 3.3.2 LIA in Western North America …...……..………...... ….10 Chapter 4: Methods….………………………...………………..……….…………….....13 4.1 Mapping………………………………...……………………………………13 4.2 Relative Dating …………………………………....………………..…….....13 4.3 10Be Cosmogenic Dating ……………………………………..………….…..15 4.3.1 Sample Collection…………….………………………………….….15 4.3.2 Laboratory Methods…………………..……..…………………..…..15 4.3.3 Determining Ages from 10Be /9Be ratios………………………..…..16 4.4 Dendrochronology/Dendroglaciology…………………..…….……..….…...17 4.4.1 Sample Collection……………………………...……...………...…..18 4.4.2 Determining a Minimum Moraine Age…………………...……...…19 4.4.3 Laboratory Methods…………………………..……………...……...20 4.4.3.1 Constructing a skeleton chronology…………….…………20 4.4.3.2 Statistical Analysis Using COFECHA……..…...…..…….21 4.5 Determining Equilibrium Line Altitudes (ELA’s)…………...... ……….….22 4.5.1 Equilibrium Line Altitude (ELA)………………..………….…...….22 4.5.2 Accumulation Area Ratio (AAR)………………………..……....….23 4.6 Climate Change Estimates…………………..……………………………….24 Chapter 5: Results……………………………………....…………..………………...….27 5.1 Mapping…………………………………………..………………………….27 5.2 Establishing Relative Moraine Ages………………………..……………..…27 5.3 10Be ages……………………………………………………………..………29 5.4 Dendrochronology…………………….………….……………………...…..29 5.5 Equilibrium Line Altitudes……………………..…...……………………….30 5.6 Climate Change Estimates………………………………..………………….31

Chapter 6: Discussion……………………………………………………..……………..31 6.1 Ages of Moraines and Striated Bedrock………………..………………...….31 6.1.1 10Be ages……………………………………………..…………..…..31 6.1.2 Dendrochronology……………..…………………………………….35 6.2 Glacier Dynamics and Growth……..…………………………………….….36 6.2.1 Glacier Response to Climate.………………..……………………....36 6.2.2 Equilibrium Line Altitude Fluctuations……..……………………….37 6.3 Climate Record and Glacier Implications………...………………...... …..38 6.3.1 Climate Signal from Dendrochronology……………..………………38 6.3.2 Climate Chronology……………………………………..………...…41 6.3.3 Climate Change Estimates……………………..…………………….43 6.3.4 Timing of Moraine Deposition………………..…………...…….…..46 6.3.4.1 M-1……………………………..………….………………46 6.3.4.2 M-2………………………..………………….……………48 6.4 Regional Significance…………………………………………………...... 51 6.4.1 Regional Perspective……………...………….……….………...……51 6.4.2 Correlation with Regional Cycles………..…………………….….…54 Conclusion…………………………………………...…………..………………………55 References…………………………………………...……………...………………..…..58 Appendix A. Figures…………………………………………………..……..….…….…66 Appendix B. Tables…………………………………………………....……...………..113 Appendix C. Accelerator Mass Spectrometer Data…………………..…………..….…119 Appendix D. COFECHA Output Data……………………………..…………...………123

List of Figures

Figure 1. Location map of the Grizzly Valley in the Klamath Mountains 66

Figure 2. DEM of Grizzly Valley cirque 68

Figure 3. Glacial maxima chronology 69

Figure 4. M-1 and M-2 moraine crests 71

Figure 5. Sample site location map 72

Figure 6. Tree species for dendrochronology 74

Figure 7. Pith correction schematic 75

Figure 8. Moraine age from dendrochronology 76

Figure 9. Skeleton plot diagram 77

Figure 10. ELA climate conditions and envelope 78

Figure 11. Location map of temperature stations 79

Figure 12. Location map of precipitation stations 80

Figure 13. M-1 and M-2 crest comparison 81

Figure 14. M-2 profile 83

Figure 15. M-1 soil development 85

Figure 16. Moraine surface soil/clast diagrams 86

Figure 17. 10Be cosmogenic ages 89

Figure 18. Grizzly Valley cirque cosmogenic sample map and ages 90

Figure 19. Tree ages with ecesis interval 92

Figure 20. Graphs of the M-1 standardized tree-ring chronologies 93

Figure 21. Tree-ring width correlations 95

Figure 22. DEM of cirque with ELAs 97

Figure 23. Climate estimates at ELAs 99

Figure 24. Location map of proximal tree-ring studies 101

Figure 25. Temperature records 102

Figure 26. Regional subdivisions for dendrochronology 104

Figure 27. Precipitation records in NW North America 105

Figure 28. Precipitation records of study region 107

Figure 29. Precipitation lapse rate 108

Figure 30. Temperature lapse rate 109

Figure 31. Tree-ring widths and ENSO record 110

Figure 32. Reconstructions of the Grizzly Valley glacier 111

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List of Tables

Table 1. Documented ecesis intervals in North America 113

Table 2. Summary of boulder and moraine characteristics 114

Table 3. Summary of M-1 soil descriptions 115

Table 4. Temperature and winter precipitation changes at calculated ELA’s 116

Table 5. COFECHA statistical analysis values 117

Table 6. Descriptions of precipitation data weather stations 118

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ABSTRACT

LATE HOLOCENE GLACIAL ADVANCES IN THE KLAMATH MOUNTAINS,

NORTHERN CALIFORNIA, DETERMINED FROM 10BE COSMOGENIC

EXPOSURE DATING AND DENDROCHRONOLOGY

Joshua Tate Graham

Master of Science in Geology

The well!preserved moraines in the cirque at the head of Grizzly Creek, Klamath

Mountains, California, provide the most complete record of late!Holocene glacier fluctuations yet documented in this region. Two separate moraine complexes lie below the modern glacier, within the cirque, only one of which supports substantial tree growth.

10Be cosmogenic ages of scoured bedrock surfaces and moraine boulders, as well as tree!ring ages indicate the approximate timing of glacial maxima in the Grizzly Valley cirque. The combination of the detailed climate record, provided by tree!ring widths, and the estimated moraine ages, determined from dendrochronology and cosmogenic dating, allows for an accurate reconstruction of the Grizzly Valley Glacier fluctuations over the last 1,000 years. Nine new cosmogenic exposure ages, combined with dendrochronology, constrain the timing of glacier maxima to ~690, ~150 and ~130 years before present

(ybp). Around 690 ybp, the equilibrium!line altitude (ELA) was depressed ~160 meters relative to the ELA of the modern glacier. Using local temperature and precipitation lapse rates and the elevation of the contemporary glacier, we found that in comparison with modern climate conditions, the mean summer temperature during the ~690 ybp glacier

ix maximum was ~0.9°C less and winter precipitation was ~95 cm in snow water equivalent

(SWE) greater. During the ~150 and ~130 ybp glacier maxima, the ELA was ~67 meters lower than the modern ELA. The mean summer temperature corresponding to this glacier maximum was ~0.4°C cooler and winter precipitation was ~44 cm greater in comparison with modern climate. The climate regime over the last 1,000 years in the Klamath

Mountain region seems to be cool and exceptionally wet. The ~690 ybp glacier maximum in the Klamath Mountains is not apparent in the Sierra Nevada or Cascade Ranges. Also, glaciers in the Sierra Nevada and Cascade Ranges retreated from their most recent LIA maxima ~20-30 years before the glaciers of the equivalent advance in the Klamath

Mountains. The climate in the Klamath Mountains likely varies from the Sierra Nevada and Cascade ranges due to the proximity to the Pacific Ocean. The chronology of Grizzly

Valley Glacier fluctuations, determined in this study, suggests that in California the regional response to large!scale climate regimes can vary over relatively short lateral distances.

Chapter 1: Introduction

The extreme climate fluctuations and associated glacial advances of the last glacial interval have received considerable attention, but there is surprisingly little known about glacial advances during the Holocene interglacial (11,500 ybp-present) (Mayewski et al.,

2004). In particular, the mid- to late-Holocene climate (last 6,000 years) is likely similar to today’s, and thus understanding this period can be useful for interpreting how present climate changes will impact our environment in the near future (Wanner et al., 2008).

With the likely human-induced global warming over the past few decades, natural climate variations during the mid- to late-Holocene provide a base line against which the extent of anthropogenic effects can be estimated (Matthews and Briffa, 2005). Despite the profound importance of understanding Holocene climate, many uncertainties remain.

Such uncertainties include whether Holocene climate fluctuations have been cyclic, as a continuation of the 1,500 year Dansgaard-Oeschger cycles (Bond et al., 2001), or better explained by regional events like changes in Pacific Ocean sea-surface temperatures due to deep-ocean circulation or tropical volcanic eruptions (Campbell et al., 1998; Wanner et al., 2008). Such debates are difficult to resolve, because the glacial record during the

Holocene is tenuous and rather ambiguous (Mayewski et al., 2004). The causes for temperature and precipitation changes during the Holocene also remain uncertain, though the answer has important implications for modern climate variations.

One way to determine late Holocene climate variations is to use glacial fluctuations during this period as a proxy for climate change. Alpine glaciers are an ideal proxy for understanding Holocene climate, because their size and shape is dependent on local climate (Thompson, 2009). Small alpine glaciers are particularly sensitive to climate

1 change, because they respond to slight changes in summer temperature and mean annual precipitation, as well as respond rapidly to natural forcing mechanisms, including solar variability, orbital variations, land cover, volcanic aerosols, and greenhouse gases (Davis, et al., 2009; Lillquist and Walker, 2006; Wanner et al., 2008). Climate forcing can produce measurable alpine glacier size changes on a scale from as few as 1-5 years and on a decadal scale (Lillquist and Walker, 2006). Therefore, alpine glacial deposits provide a high-resolution record of climate fluctuations that other proxies, such as ocean sediment cores, pollen records, and ice cores, may not recognize (Marcott, 2005). These paleoclimate records can also be used to test effects of synoptic-scale forcing mechanisms including El Niño-Southern Oscillation, Pacific Decadal Oscillation, and regional atmospheric circulation patterns (Bowerman and Clark, 2005).

Understanding alpine glacier variations and associated climatic events is not only critical for climate change assessment, but also for resource management. Changes in glacier mass significantly affect downstream water flow, which greatly impacts energy production, water resources, and agriculture (Davis et al., 2009; Koch et al., 2009;

Tangborn, 1980). Snowpack in northern California has a dramatic influence on the water supply for the entire state, which accounts for 10% of the nation’s population. Foreseeing future climate trends, which is largely done by analyzing late Holocene climate fluctuations, is invaluable to society, enabling us to allocate limited resources in the most efficient way possible.

Glaciers in the Trinity Alps have left an exceptionally well-preserved depositional record of their movements since the Last Glacial Maximum (~15,000 years), including well-defined late-Holocene moraines. Here, we date the remarkable late-Holocene

2 depositional record of the Grizzly Glacier using high-precision 10Be cosmogenic nuclide exposure techniques and dendrochronology to reconstruct regional climate patterns in northern California. Because chronologies of glacial maxima over the last millenia, including the Little Ice Age (LIA) advances, are variable in the Sierra Nevada, Cascades, and Coast Mountains, this study investigates whether a climatic relationship exists between the Trinity Alps region and these mountain ranges during the late-Holocene or if the Trinity Alps have an independent climate regime. With moraine ages and reconstructed equilibrium line altitudes (ELA) of former glacier positions, both of which are calculated in this study, it is possible to determine the approximate temperature and precipitation differences, relative to today, during the deposition of each dated moraine.

Quantifying temperature and precipitation variations in the late-Holocene is essential for understanding the magnitude of climate fluctuations in the recent past and comparing those anomalies with the rate and magnitude of change today. Dendrochronology in this study establishes a detailed record of climate fluctuations over much of the last half millennia, which is compared with other reconstructed climate chronologies to further assess whether a regional synchronicity exists in climate patterns between northern

California and other glaciated (or once glaciated) regions of western North America. The

Grizzly Glacier moraines offer a rare opportunity to carry out a multi-proxy (10Be cosmogenic and dendrochronology) approach to reconstruct the scarcely documented late-Holocene climate patterns in northern California and contribute to the nebulous glacier chronology of western North America.

Chapter 2: Study Area

Thompson Peak is located in the western-central part of the Trinity Alps

Wilderness (41.0005 N, 123.048 W), which are in the southern part of the Klamath

Mountains (Figure 1). Thompson Peak is the highest peak in the Trinity Alps (2741 m).

This region was heavily glaciated during the Pleistocene, and a few small glaciers and snowfields remain in sheltered cirques today (Sharp, 1960). These contemporary glaciers

(~2500 m) subsist more than 610 meters below the elevation of modern glaciers in the

Sierra Nevada due to the more northerly latitude and nearness to the Pacific Ocean

(Guyton, 1998). The Grizzly Valley cirque and glacier are located on the northern side of

Thompson Peak (Figure 2). The relatively high elevation and topographic shading from

Thompson Peak results in enhanced snow accumulation on the north side of the peak and allows the unnamed glacier (termed the Grizzly Valley Glacier for the remainder of this paper) to persist through present time, while most glaciers in the western United States have been receding rapidly (Howat et al., 2007). The cirque is fixed in the Canyon Creek

Pluton, a Jurassic/Cretaceous tonalite batholith, which intruded into Permian meta-sedimentary marine rocks (Allen and Barnes, 2006). The glacial deposits studied in this investigation were all derived from the tonalite batholith.

The present climate of the study area is Mediterranean (Skinner, 2003), characterized by cool, wet winters and warm, dry summers (Skinner et al., 2006).

However, due to a strong moisture and temperature gradient from the west to east, resulting from the nearness to the Pacific Ocean and significant elevation variations, local expression of this climate regime is highly variable (Skinner et al., 2006). Typically, most precipitation falls between October and April as snow and the amount of precipitation

4 decreases with distance from the ocean. More precipitation generally falls in areas of high elevation and significant relief over short lateral distances (Skinner et al., 2006), consistent with the Grizzly Valley cirque. Therefore, this particular cirque is in an ideal location for accumulating anomalously high snowpack each year.

Chapter 3: Background

3.1 Glacier Response to Climate

The elevation and position of glacier termini are generally thought to provide the most reliable parameters for quantifying glacier response to climate (Nesje and Dahl,

2000). The advance and retreat of glacier termini and fluctuations in ice volume are a direct function of ice accumulation and ablation (Bahr et al., 1997). When accumulation is greater than ablation, glaciers are under positive mass balance, forcing them to thicken and advance (Lillquist and Walker, 2006; Roe and O'Neal, 2009). Conversely, glaciers thin and retreat when summer ablation, due to surface energy exchanges, exceeds winter accumulation (Burbank, 1982; Lewis and Smith, 2004). In northwest North America, winter accumulation anomalies are primarily attributed to precipitation, rather than temperature (Bitz and Battisti, 1999; Burbank, 1982; Lewis and Smith, 2004). However, during the ablation-season, temperature anomalies are the dominant influence on the glacier mass balance, and thus, on glacier response to climate fluctuations (Burbank,

1982). For the northern Cascades, studies show that a decrease in summer temperature of just 0.5 degrees or an increase in winter precipitation of 10% is sufficient to change glacier mass balance (Tangborn, 1980), resulting in glacial advance. It is important to recognize that the mass balances of glaciers within a particular region often co-vary

(McCabe et al., 2000). Although glaciers generally advance and retreat due to changes in climate, glacier size is dependent on ice accumulation and ablation, which are a function of several environmental factors including: altitude, avalanches or wind-blown snow, valley orientation, rain versus snow, surface albedo variations, valley geometry, hillside shading, debris cover and cloudiness (Rupper and Roe, 2007). These local conditions can make glacier response to climate highly variable over short distances. However, when numerous glaciers over a larger spatial area advance and retreat correlatively, it can be assumed that regional-scale climate changes are dominating glacier behavior and local factors are not controlling glacier response (Rupper and Roe, 2007).

3.2 Holocene Climate Variability

Climate during the Holocene was once considered to be fairly stable, especially in comparison with the climate variations of the Last Glacial Maximum (LGM). Continued investigations have caused researchers to acknowledge multiple cooling events during the present interglacial period, including the Little Ice Age (~700-150 ybp). Following the

Younger Dryas (~12.9-11.7 k.y), a period of warm post-glacial temperatures commenced, termed the Holocene climatic optimum (~9-6 k.y.) (Rossignol-Strick, 1999). After the

Holocene climatic optimum, western North America entered a phase of gradual cooling.

The long-term cooling initiated multiple alpine glacial expansions over the past ~5 k.y., termed Neoglaciation (Denton and Karlen, 1973; Konrad and Clark, 1998; Bowerman and Clark, 2005). This cooling was interrupted when North America, Europe, Greenland, and Iceland experienced a relatively warm climate interval, known as the Medieval

Warm Period (MWP), from 1100-810 ybp (Hughes and Diaz, 1994). The cooling during the Neoglacial period culminated in the Little Ice Age (LIA) (Schimmelpfennig et al.,

2012), which is widely recognized as the most extensive glacial advance during the

Holocene (710-110 ybp) (Davis et al., 2009; Matthews and Briffa, 2005). In northern

California, the most extensive LIA glacial maximum occurred around 250-150 ybp, evident in the Sierra Nevada and Cascade Ranges (Bowerman and Clark, 2005; Sigafoos and Hendricks, 1972; Koch et al., 2009). Glaciers in the European Alps, Iceland and

Scandinavia also reached a maximum extent from ~700-600 ybp (Davis et al., 2009).

Although the Milankovitch orbital cycles, with periodicities from 19,000-308,000 years, exhibit very consistent frequencies and coincide with the Last Glacial Maximum

(~18,000 ybp), they do not explain postglacial short-term (i.e. millennial) climate fluctuations (Campbell et al., 1998). The Younger Dryas (~12,000 ybp), Neoglacial advances (5,000-3,300 y.b.p.), Medieval Warm Period (1,200-700 ybp), and the LIA

(600-140 ybp) have been attributed to millennial-scale patterns with ~1500 year periodicities, commonly referred to as Dansgaard-Oeschger cycles (Campbell et al.,

1998; Bond et al., 2001; Solomina et al., 2008). The global nature of these 1500 yr climate frequencies suggests a large-scale forcing mechanism, such as cyclic fluctuations in solar output or changes in deep-ocean circulation (Campbell et al., 1998; Bond et al.,

2001; Solomina et al., 2008). Denton and Karlen (1973) recognized another set of late

Holocene millennial-scale cycles with a ~2500 yr periodicity, which also appear to fit the timing of the LIA (Porter, 1986). These 2500 year climate pattern frequencies are evident in time series reconstructions derived from oxygen isotope investigations from the

Greenland and Devon Island ice cores (Fisher, 1982), as well as during the last 8500 years in the Camp Century Greenland ice core (Dansgaard et al., 1984). However, the forcing mechanism for these 2500 climate patterns remains undefined (Porter, 1986).

Evaluating the global synchronicity of decadal to multi-decadal scale glacial fluctuations is more challenging, due to the limited resolution of most climate and glacier proxies, as well as local climate forcing mechanisms that occur at a decadal to multi-decadal scale, such as the Pacific Decadal Oscillation and El Niño-Southern Oscillation (Solomina et al., 2008).

Licciardi et al. (2009) propose that due to a decrease in Gulf Stream current during the time of the LIA, the oceanic heat transfer to the north was cut off. Subsequently, cool waters north of the equator caused a southward shift of the Intertropical Convergence

Zone (ITCZ) leading to anomalously cool conditions to the north, favorable for glacier growth (Licciardi et al., 2009). Although the forcing mechanism for global decadal- to century-scale climate oscillations, including the multiple stages of the LIA, is largely unknown, regional variations are known to be influenced by cyclic phenomena including the North Atlantic Oscillation (NAO), El Niño-Southern Oscillation (ENSO), and Pacific

Decadal Oscillation (PDO) (Solomina et al., 2008; Papineau, 2001). These relatively small-scale climate oscillations shifted oceanic temperature gradients and ocean circulation, causing changes in sea surface temperatures (SST). These sea surface temperature variations affected mean annual temperature and precipitation across the

Pacific Ocean, and consequently the Pacific Northwest. (Papineau, 2001).

Various external, non-cyclic forcing mechanisms, such as solar variability, sulfate aerosols from volcanic eruptions, and increased flux of atmospheric CO2, may also have contributed to late-Holocene climate variations, including the Little Ice Age, (Davis et al.,

2009; Lillquist and Walker, 2006; Porter, 1986; Wanner et al., 2008). The Greenland ice core shows a strong connection between increased ocean acidity and Northern

Hemisphere glacier fluctuations on a decadal-scale, suggesting the induction of sulfate aerosols into the atmosphere from volcanism has a significant influence on global decadal-scale climate variations, sufficient to induce glacial advance during the LIA

(~0.5-1.2oC) (Porter, 1986). Realistically, some combination of these proposed forcing mechanisms are likely responsible for late-Holocene glacier-climate fluctuations.

However, high resolution chronologies of glacial fluctuations, from different locations during the LIA, are required to test models of late-Holocene climate forcing (Licciardi et al., 2009).

3.3 Previous Studies

3.3.1 LIA in the Northern Hemisphere

Some evidence suggests that the LIA was a global phenomenon (Kreutz et al.,

1997), but difficulties in determining precise ages of glacial advances have made it challenging to define convincing relationships between localities (Licciardi et al., 2009).

However, the LIA does appear to be at least hemispheric in extent, beginning around 760 y.b.p. (Schimmelpfennig et al., 2012). Porter (1981) acknowledged a synchronicity in

Northern Hemisphere glaciers during the LIA, distinctive from Southern Hemisphere variations, likely due to dissimilar volcanic activity in the two hemispheres. Glacial fluctuations are highly synchronous in the Western Alps in Europe. Recent studies show these glacier fluctuations have a strong correlation with temperature changes in the North

Atlantic and the tropical hydrological cycle, which supports the idea of a hemispherical climate link during the Holocene (Schimmelpfennig et al., 2012).

The Northern Hemisphere generally had two main phases of the LIA, the first occurring from 760-530 ybp and a second, main phase, from 460-110 ybp (Porter, 1986).

The first phase is well documented in reconstructed glacial chronologies of the, North

America, Scandinavia, Greenland, Iceland and the Himalayas (Grove, 1988). Climate reconstructions based on tree-ring analysis in almost all areas of the Northern

Hemisphere reveal a significant decrease in summer temperatures (0.0-2.0°C below the

1961-1990 mean) from around 440-110 ybp (Matthews and Briffa, 2005). Not surprisingly, this temperature depression coincides with the second, main phase of the

LIA, which is documented in nearly all glaciated regions of the Northern Hemisphere.

Between these two advancing stages, most glaciers were in a recessive state, but not to the extent of conditions during the Medieval Warm Period (Porter, 1986). In Europe and

North America, glacier chronologies of the main phase (460-110 ybp) of the LIA suggest that glaciers at higher latitudes advanced late and retreated early, while glaciers at lower latitudes advanced relatively early and re-advanced multiple times to comparable positions, before receding significantly in the first half of the twentieth century (Grove,

1988).

3.3.2 LIA in Western North America

Previous studies have contributed to a continuously developing Holocene glacier chronology of western North America (Figure 3), which remains rather incomplete and poorly understood. A gap of ~13,000 years exists in the glacial record of the Sierra Nevada, between the Recess Peak (~13 ka) and Matthes (<200 ybp) moraines

(Bowerman and Clark, 2005; Gillespie and Zehfuss, 2004). Evidence from lake cores, and the absence of moraines between the Recess Peak and Matthes moraines, indicate that the climate of the Sierra Nevada during the LIA was possibly the coldest and wettest of the last 13,000 years (Gillespie and Zehfuss, 2004). The only convincing dates to

10 dispute this claim were presented by Konrad and Clark (1998), which indicate an advance around 3400 ybp, derived from coring Conness Lake. Their cores also yielded ages synchronous with other documented Matthes advances between 100 and 700 ybp

(Konrad and Clark, 1998). Bowerman and Clark (2005) interpreted two distinct glacial maxima over the last millennia at ~700 and ~250-170 ybp, the most recent being the most extensive Little Ice Age advance (LIA maximum). Scuderi (1987), based on dendrochronology, interpreted three maxima in the Sierra Nevada, over the last 1,000 years, as distinct stages of the Matthes advance, occurring around 700-600, 450-300, and

200-150 ybp.

In the Cascade Range, the Coleman and Deming glaciers on Mount Baker advanced as early as 990 ybp and reached their most extensive maximum around 560-390 ybp

(Osborn et al., 2012). These dates are synchronous with the maximum LIA ages determined by Davis et al. (2009) on Mt. Baker around 580–380 y.b.p. The mid- sixteenth century advance was also documented by dating moraines of the Southern

Cascade and LeConte glaciers (Grove, 1988). After this advance, glaciers in the Cascades generally receded before readvancing to less extensive positions in the late eighteenth to early nineteenth century (Grove, 1988). Lichenometric investigations of moraines on Mt.

Rainier, Washington, indicate glaciers here reached a maximum extent around 230-200 ybp (Burbank, 1981), which may have been slightly earlier than other glaciers in the

Cascades, due to the relatively high elevation (4392 m) of this volcano. Based on tree-ring ages, eight Mt. Rainier glaciers started to recede synchronously around 150-170 ybp (Sigafoos and Hendricks, 1972). Late eighteenth century to mid nineteenth century maximums were also determined for the Deming and Boulder Glaciers on Mount Baker,

Washington, in the Cascade Range (Burke, 1972; Easterbrook and Burke, 1971; Fuller,

1980), which were in agreement with the 160 ybp advance of the Blue and Hoh glaciers on Mt. Olympus, in western Washington (Heusser, 1957). Therefore, the major advances over the last millenia in the Cascades appear to have occurred from the mid fifteenth century to early seventeenth century, and late eighteenth century to mid nineteenth century, after which temperatures during the ablation season increased and precipitation decreased.

In Alaska, the Juneau Icefield, a heavily glaciated region of the Coast Range, had synchronous glacial advances around 250-210 ybp and ~135 ybp (Grove, 1988).

Evidence from the Klinaklini Glacier, in the southern Coast Mountains, indicates a glacier advance initiated around 1020-1080 ybp, interpreted as a relatively early start of the LIA, which reached its Holocene maximum extent between 280-150 ybp (Menounos et al., 2009), synchronous with the glaciers of the Juneau Icefield. The glaciers of

Garibaldi Provincial Park, British Columbia, also located in the southern Coast

Mountains, synchronously reached a LIA maximum around 320-290 ybp and several less extensive advances occurred from 200-90 ybp (Koch et al., 2009). The Taku glacier, one of the largest glaciers in the Juneau Icefield, readvanced from 1890 to 1952 at a rate of

100 m per year (Grove, 1988). The glaciers in the Coast Mountains have a rather variable chronology, probably due to the high relief and proximity to the Aleutian Low atmospheric current.

Chapter 4: Methods

Mapping and sample collection took place over an 8 day period in August 2012.

Undergraduate students (during field work in August 2012) Jose Cardona and Ryan

Witkosky from the Department of Geological Sciences at California State University,

Northridge assisted with the fieldwork in the Klamath Mountains. Dr. Richard

Heermance also assisted with the field and laboratory work.

4.1 Mapping

The moraines of late-Holocene Grizzly Lake Glacier maxima were distinguished and mapped in detail within the cirque beneath the Grizzly Valley Glacier. Aerial photographs and satellite imagery were used in the field to supplement mapping. GPS coordinates were also used to constrain moraine features and sample locations. Two moraines of different ages (M-1 and M-2) were distinguished in the field based on stratigraphic position, extent of lichen growth on boulders, soil development, surface weathering, extent of vegetation, and moraine preservation.

4.2 Relative Dating

Various field methods and carefully documented observations served as indicators of relative glacial landform ages. The most obvious field observation for determining relative ages of M-1 and M-2 was the position of the landforms relative to one another in the cirque; down-valley moraines are older than those deposited up-valley. Subsequent glacial advances of greater magnitude, extending further down-valley, would destroy any previously deposited moraines of smaller advances. In order to determine a soil-to-clast ratio on each moraine, as well as to provide a mean clast size for the crest of each moraine, clast counts were taken along 50-meter transects of moraine crests (Figure 4). A

13 higher clast-to-soil ratio is generally more typical of a younger moraine, because soil has not had time to develop. The degree of mineral weathering/degradation and discoloration on moraine boulder surfaces also indicated relative moraine ages, with these characteristics becoming more pronounced with age. The extent of lichen growth on moraine boulders, which varied greatly between M-1 and M-2, also provided a proxy for moraine age. The degree of boulder stability was also assessed as an indicator of moraine age; loose, unstable slopes are characteristic of recent moraine deposition, while well-consolidated slopes are characteristic of older moraines (Bowerman and Clark,

2005). The angularity of the boulders was also an important field observation, due to the relationship between moraine age and boulder roundness. Typically, fresh, angular, unweathered boulders suggest a recent moraine deposition (<200 ybp for this investigation). Soil development and associated vegetation cover were also documented and provided additional insight for determining a relative moraine age.

The degree of moraine degradation was the most apparent distinguishing characteristic between the moraines in the cirque. Moraine profiles were created in the field using a clinometer and tape measurer and were later digitized. The profiles are indicative of moraine age through the relationship between time and moraine degradation; moraine slope angles decrease over time, transforming from sharp to broad crested. Moraine degradation is most vigorous immediately after a glacier recedes from the moraine surface, on a decadal to centennial scale (Putkonen and Michael, 2006).

In summary, moraines that had rounded boulders with extensive lichen growth, and weathered surfaces with rounded crests and stable slopes were considered older than those with more angular boulders, less lichen cover, less weathered clasts, and steep,

14 unstable slopes with little soil development. The degree of vegetation cover was also used to indicate the relative age of the moraine deposits, because development of vegetation increases with time since moraine deposition.

4.3 10Be Cosmogenic Dating

4.3.1 Sample Collection

We collected 10 surface samples from large moraine boulders on moraine crests and analyzed 7 samples to determine exposure ages (Figure 5). Samples were collected from boulders greater than 1 m in height to diminish the shielding effects of exhumation and snow cover. At each boulder, we sampled the upper 0-5cm of bedrock with rock hammers and chisels. At least a kilogram of rock from each sampled boulder was collected and stored in a bag labeled with the appropriate sample name. At each sample site a photograph and GPS location were taken and the following were documented: thickness of sample taken from boulder, height of the sample from ground, and boulder dimensions. Topographic shielding was documented by measuring the inclination or the topographic horizon from each sample location. The sampled rock surface was described

(whether it was a point, flat surface crest, protruding felsic vein, etc.) and the trend and plunge of the sampled boulder face recorded.

4.3.2 Laboratory Methods

All Grizzly Lake moraine sample processing was conducted in the rock lab and clean lab at California State University, Northridge. Rock samples were crushed with a jaw crusher into gravel (<10mm) and then ground with a Bico Grinder to ~1 mm. All samples were then dry-sieved to separate <0.250 mm, 0.250 – 0.600 mm, and >0.600 mm grain sizes fractions. For each sample, approximately 500 grams of the 0.250 – 0.600 mm

15 grain size fraction was used to continue sample processing. All samples were processed in accordance with the standard techniques described by Kohl and Nishiizumi (1992) and refined by Bodo Bookhagan at the University of California Santa Barbara Cosmogenic

Nuclide Preparation Facility. A Frantz magnetic barrier separator was used to carry out the magnetic mineral separation. Organic material and carbonates were removed from the samples in an 8 hour 6N hydrochloric acid (HCl) bath at ~70 °C. Quartz purification and etching was furthered by subjecting the samples to a 6-hour 2% hydrofluoric acid

(HF)/2% nitric acid (HNO3) bath on bottle rollers at ~65 °C. This step was followed by three 6-hour steps of 1%HF/1%HNO3 at ~65 °C. After the samples were purified, they were dissolved in 49% HF and washed 4 times with a 3:1 HCl:HNO3 solution (aqua regia). A Be carrier prepared in-house by Richard Heermance was then added to each sample and the blank. The in-house carrier was made from a phenacite crystal from the

Ural Mountains with a 10Be /9Be ratio of 5 x 10-15. Ion-exchange column separation chemistry was then used to isolate the Be from the samples. The Be was then precipitated as BeOH from a solution by adding NH4OH and incubating the sample overnight. The

BeOH was dried down and combusted in boron-free glass tubes. The concentrated BeOH residue was then packed into AMS targets with niobium powder for AMS analysis.

4.3.3 Determining Ages from 10Be /9Be Ratios

The 10Be concentration in the rock sample is a direct result of the time period the boulders have been exposed at the surface. Once the boulders are deposited at the surface, exposure to cosmogenic radionuclides causes a reaction with the silica and oxygen in the quartz, resulting in 10Be accumulation in the quartz lattice. 10Be production rates, scaled for given latitude and altitudes, have been well-constrained and calibrated in northeastern

North America by Balco et al. (2009). Therefore, when erosion and radioactive decay are accounted for, the exposure age of a quartz-bearing rock can be calculated based on the measured cosmogenic radionuclide concentration and the calculated production rate.

Moraine ages were calculated using the CRONUS-Earth online calculator version 2.1

(Balco et al., 2008). Although erosion corrections were applied to the data, samples collected from M-1 and M-2 were deposited at the surface so recently that exposure age error from erosion is negligible. Samples collected from striated bedrock have also undergone insignificant erosion since the ice contraction, evident from the preserved striations, so an erosion rate of 0 was assumed.

4.4 Dendrochronology/Dendroglaciology

Dendroglaciology is a subdivision of dendrochronology that uses tree ring analysis to better understand and date the movement of glaciers (Smith and Lewis, 2007).

Dendroglaciology is an effective proxy for determining moraine ages and interpreting subsequent climate conditions, because trees only grow on moraines after glacier ice has receded from the moraine surface and grow thereafter at variable annual growth rates in response to climate conditions. Dendrochronology is particularly suited for dating late

Holocene landforms in comparison with other geobotanical methods, because it provides ages with annual resolve accuracy (Smith and Lewis, 2007).

In this study, cores were taken from two tree species growing on the surface of the moraines: Foxtail Pine (Pinus balfouriana) and Mountain Hemlock (Tsuga mertensiana)

(Figure 6). Foxtail Pines typically inhabit rocky slopes above 1525 m and only grow up to about 14 m tall (Jones, 1986). Mountain Hemlocks are characteristically found on open slopes at higher elevations (1600-3000 m) and reach heights of up to 30.5 m (Jones,

1986). Variations in weather and climate generate a physiological response in trees, directly impacting their phenology, including annual tree growth rate (Lo et al., 2010).

Glacier mass balance is also dependent on winter precipitation and summer temperature.

Therefore, an inherent relationship exists between radial growth (ring width) in climate sensitive trees and glacier mass balance (Smith and Lewis, 2007).

4.4.1 Sample Collection

A total of 39 cores were collected in the Grizzly Valley cirque from living trees growing on the surface of moraines. The number of sample collection sites on each moraine can be summarized as follows: 9 cores on M-2, 10 cores on the east limb of M-1,

12 cores on the west limb of M-1, and 8 cores on the western moraine sequence (Figure

5). Tree core samples were only taken from living trees with a large circumference relative to the rest of the tree population on the same moraine, in order to sample the oldest tree on each moraine. Although the largest tree on the moraine is not always the oldest tree on the moraine, by sampling a collection of the largest trees growing on the moraine, the oldest tree will likely by sampled. Due to the limited number of trees growing on M-2, because of the very recent moraine deposition, all trees with a large enough diameter to be cored were sampled. Tree cores were taken with an increment borer and stored in paper straws with an arrow indicating the direction of the most recent tree ring, which typically still had bark attached. Cores were taken as close to the ground as possible to obtain the most accurate tree-ring count. Multiple observations and field notes were recorded at each tree sampled: the tree’s diameter at breast height (DBH), height sampled from ground, GPS location, picture number, tree species and any unique tree characteristics. Pictures were taken of each tree sampled with a field assistant or

18 backpack for scale.

When removing tree cores in the field, it was not always possible to recover the pith of the tree (center of the tree). Circumstances including asymmetric radial growth, low tree branches, and difficulties in aiming the corer exactly into the center of older, larger trees occasionally resulted in off-pith sampling, which required the addition of a correction factor to account for all radial rings. The pith correction factor (PC) was based on the curvature and mean width of the innermost rings (Figure 7), which was added to the original ring count of the core. Therefore, a simple equation was used to calculate the age of moraine deposition, summarized in Figure 8.

4.4.2 Determining a Minimum Moraine Age

The age of the oldest tree growing on the surface of the moraine, determined by counting the annual growth rings, provides an estimate of the minimum age for the moraine (Zhu et al., 2012). Tree-ring counting yields a limiting age for the moraine because the tree can only begin to grow after the moraine surface has stabilized and formed soil, following glacial retreat (Koch et al., 2009). However, to establish a more accurate minimum moraine age, the ecesis interval (lag time between glacier retreat and tree germination) is added to the age determined from counting annual tree rings

(McCarthy and Luckman, 1993). Previous investigations have determined ecesis intervals for numerous tree species in the North American Cordillera, which range from 1 to 96 years (Barclay et al., 2009; Lewis and Smith, 2004; Winchester and Harrison, 2000;

Winchester, 2001; Luckman, 1988; Ruiz et al., 2011; Smith et al., 1995; McCarthy et al.,

1991; McCarthy and Luckman, 1993), summarized in Table 1. The ecesis interval (EI) is a function of multiple factors, including the existence of a seed source, seedbed

19 characteristics, and climatic conditions during the ecesis interval (Smith and Lewis,

2007). In the Grizzly Valley cirque, seeds are readily available down-valley and would only be deposited by birds or strong up-valley wind flow. The climate conditions in the cirque are rather harsh for a seedling and the resistant granite soil source is not particularly supportive of seed germination. Therefore, the ecesis interval for the trees growing on the M-1 and M-2 moraines is most likely between the mean ecesis interval documented in North America and the maximum interval documented (32-96 yrs).

4.4.3 Laboratory Methods

4.4.3.1 Constructing the Skeleton Plot Chronology

After removing the tree cores from the paper transporting straws and visually inspecting them, all cores with significant amounts of rot were disregarded for analysis

(9). All cores suitable for analysis (30) were glued into wooden mounts and sanded with an orbital sander using progressively finer sandpaper of the following grits: 120, 220,

320, and 400. Once each core was sanded down to about half of its original diameter and the flat surface was well–polished, the annual tree rings were ready for visual analysis.

The tree rings were first dated using the skeleton plotting technique. Skeleton plots were constructed for all cores in accordance with standard techniques described by Speer

(2010). The plots were made on graph paper with five squares to a centimeter, each vertical line representing one annual year. The annual vertical lines form the basis on which the cores can be compared to one another. The length of each line correlates with the width of the tree ring, ranging from 0 squares (for relatively wide rings) to 10 (for relatively narrow rings), as seen in Figure 9. Therefore, the skeleton plots can be visualized as a 2-D plot with the x-axis representing time, going from present on the right

20 to older on the left, and the y-axis representing an inverse scale of the narrowness of the ring (Speer, 2010). The trees were then compared with each other to identify marker rings, or years with consistently narrow (long line) or wide (short line) rings. Because trees tend to have wider rings in the early years and progressively narrower rings towards the outside of the tree, a mental standardization process must be applied while constructing the skeleton plots (Stokes and Smiley, 1968). This standardization process is done by comparing the ring being dated with the width of the four rings on either side of it, creating a dynamic calibration for ring widths. After completing each skeleton plot, the rings were totaled and recorded as the ring count (RC). Once skeleton plots were made for all cores, a master chronology was built. The master chronology indicates all marker rings that consistently agree between cores. In order to make the master chronology, individual skeleton plots were overlapped, precisely lining up corresponding marker rings. Lines were drawn on the master chronology for each year if they were representative of at least 50% of the individual skeleton plots (Stokes and Smiley, 1968).

The length of the vertical lines was based on a rough average of all rings for that particular year.

4.4.3.2 Statistical Analysis Using COFECHA

Using a Velmex measuring system and microscope, annual rings in all cores were measured to 0.001 mm precision. Each core was measured in a statistical category according to the moraine it was sampled from and tree species. Foxtail Pines and

Mountain Hemlocks vary in annual growth rate, therefore statistical analysis of tree ring widths from the two species would likely provide erroneous results with low series intercorrelation. Trees from the east and west limbs of M-1 were also analyzed in

21 individual categories due to variability in steepness and sunlight exposure. Annual tree-ring widths measured with the Velmex measuring system were stored in the J2X software program, where each ring measurement could be viewed and/or modified. The measurements were then imported into COFECHA, which carries out the statistical analysis for cores within each category of moraine. COFECHA standardizes each core with a 32-year cubic smoothing spline of the tree ring widths and creates a master chronology by averaging the index values for each core (Speer, 2010). The master chronology is useful for identifying missing years in certain cores and interpreting climate conditions during tree growth. If trends of positive and negative deviations from mean correlate between the ring width master chronology and instrumental river discharge data, it can be assumed that trends throughout the entire master chronology reflect precipitation conditions. Comparing the master chronology from this study to those of nearby dendrochronology investigations also helps determine the validity of our results.

4.5 Determining Equilibrium Line Altitudes (ELA’s)

4.5.1 Equilibrium Line Altitude (ELA)

The equilibrium line altitude (ELA) is a theoretical line on a glacier indicating the altitude where annual accumulation of snow equals ablation and the glacier mass balance equals zero (Aa, 1996). Therefore, the ELA divides a glacier into zones of net mass gain

(accumulation zone) and net mass loss (ablation zone) (Osipov, 2004). ELA fluctuations over time are driven by changes in climate, primarily summer temperature and winter precipitation (Bowerman and Clark, 2005). Thus, reconstructing former ELA’s is essential for effective paleoclimate reconstruction. Modern ELA’s can be calculated

22 based on direct mass balance measurements on the glacier, while determining paleo-ELA’s require more indirect approaches (Osipov, 2004). These indirect approaches include both the area accumulation ratio (AAR) and the maximum elevation of lateral moraines (MELM) methods, which have been considered reliable for determining paleo-

ELA’s for alpine and cirque glaciers (Meier and Post, 1962; Tornes et al., 1993). The uppermost section of the east limbs of M-1 and M-2 buttress against steep, exposed bedrock where moraine preservation is lost and it is not possible to definitively identify the maximum elevation of the moraine. Therefore, we use the accumulation area ratio

(AAR) method to calculate equilibrium line altitudes of the Grizzly Glacier during the deposition of M-1, M-2, and the modern glacier.

4.5.2 Accumulation Area Ratio (AAR)

The accumulation area ratio (AAR) method is defined as the ratio of the accumulation area to the area of the total glacier when it is in steady state equilibrium

(Aa, 1996). The AAR is a reliable method for calculating ELA’s because it is directly related to the glacier-specific net budget and surface area (Meier and Post, 1962; Osipov,

2004), meaning the ELA calculated with the AAR method fluctuates as a function of glacier mass balance (Aa, 1996). A typical AAR for a cirque glacier in a climatic and dynamic steady state has an accumulation area to total area ratio of 0.65+/- 0.05 (Dahl et al., 2003). In this study, the outer limits of the glacier were reconstructed from the marginal moraines (M-1 and M-2), which were digitized in Google Earth Pro to find the total area of the glacier at the time of moraine deposition. Then the accumulation zone boundaries for the glacier during the time of M-1 and M-2 deposition were outlined so that the area of the accumulation zone corresponded with the total glacier area as a ratio

23 of exactly 0.65:l. The resulting altitude of the accumulation-ablation zone boundary then defined the ELA of the glacier at the respective glacial extents (M-1 and M-2). Based on comparing the modern glacier size with aerial photographs from 1952, the Grizzly

Glacier appears to be roughly the same size, and thus in steady state equilibrium.

Therefore, the ELA for the modern glacier was calculated with the same procedure as the reconstructed glaciers (M-1 and M-2), with an AAR=0.65:1.

4.6 Climate Change Estimates

Because a modern glacier exists in the study area, it is possible to quantify the equilibrium line depression between the modern ELA and reconstructed ELAs, as well as to compare the current climate conditions at modern and paleo-equilibrium lines

(Leonard, 1989). Previous research indicates that winter precipitation and summer temperature account for approximately 90% of snowline altitude changes (Porter, 1977).

Leonard (1989) developed a range of ELA climate conditions (summer temperature and winter precipitation), based on equilibrium line altitude climate conditions of 32 late

Pleistocene glaciers in the Colorado Rocky Mountains, South Cascades, Soviet Central

Asia, and Norway (Figure 10). The climate conditions plotted on Figure 10 show a relationship between summer temperature and winter accumulation for glacier ELAs, which can be constrained within an envelope of climate conditions. This envelope of climate conditions can be arithmetically defined through two equations developed by

Kotlyakov and Krenke (1982):

Equation 1.

2.85 Aw = 1.33(Ts + 6.66) 2.85 Aw = 1.33(Ts + 9.66)

Equation 1. Aw is water equivalent of winter accumulation in millimeters and Ts is mean summer temperature (June-August). The two equations define the left and right margins of the climate envelope in Figure 10, respectively.

By plotting the current climate conditions at the reconstructed ELAs on the climate envelope, created by Leonard (1989), it is possible to approximate the summer temperature and winter precipitation change necessary to advance the modern glacier to its M-1 and M-2 extent. To determine contemporary climate conditions at the modern and reconstructed ELAs in the Grizzly Valley cirque, historical climate data from weather stations proximal to the study area were used to determine local temperature and precipitation lapse rates (rate of change with altitude). The results were then extrapolated to the altitudes of the modern and paleo-equilibrium lines. As discussed above, mean summer temperature (June-September) and mean winter precipitation (November-March) have the greatest influence on the mass balance of mid-latitude glaciers (Leonard, 1989).

To determine the local temperature lapse rate, mean June-September temperatures

(1971-2000) were plotted from nine nearby weather stations: Ship Mountain, Big Bar 4

E, Trinity River Hatchery, Weaverville, Cecilville, Yreka, Callahan, Weed Fire Dept., and Camp Six (Figure 11) at their respective elevations and fit a linear regression to the data. The weather stations were selected based on altitude and their nearness to the study site in order to create a plot of temperatures at a wide range of elevations. Historical climate data was accessed through the California Department of Water Resources website

(http://www.water.ca.gov/floodmgmt/hafoo/csc/climate_data/northcoast.cfm).

Determining the local precipitation lapse rate is more complex than temperature due to the effects of local topography and variable moisture supply in mountainous regions (Leonard, 1989). The snow water equivalent of late season (April 1) snowpack is an accurate proxy for winter net balance on the surface of a glacier, because almost all winter snow accumulation remains (summer melt has not ensued) (Leonard, 1989).

Maximum annual snow water equivalent measurements from four local weather stations were averaged from 1947-1957. Precipitation data from this time period (1947-1957) was used because it provided the only period where precipitation data was recorded simultaneously between weather stations at high and low altitudes proximal to the study area. Because precipitation is significantly influenced by the intense rain shadow effect in this region (Skinner and Taylor, 2006), weather stations chosen for local lapse rate construction are all located 9-16 km east of the Grizzly Valley cirque (Figure 12). The historical climate data from the local weather stations (Red Rock Mountain, Bear Basin,

Big Flat, and Whalen) was accessed through the California Department of Water

Resources website. The snow water equivalent (SWE) data were used to create a power regression curve representing local precipitation lapse rate, which is a considered a more accurate depiction of the precipitation-elevation relationship than a linear regression line

(Bowerman and Clark, 2005). This power regression curve is propagated to allow estimation of precipitation patterns at the modern and paleo-equilibrium lines (M-1 and

M-2) of the Grizzly Valley Glacier.

Chapter 5: Results

5.1 Mapping

Mapping the geomorphic landforms in the Grizzly Valley cirque reveals easily recognizable, distinct moraine deposits, which provide the basis for establishing a chronology for the late-Holocene fluctuations of the Grizzly Valley Glacier. The boundary of the M-2 moraine structure, based on till accumulation, is well defined and the feature is clearly a composite moraine. The M-2 moraine consists of two closely spaced crests (< 22 m apart) at the terminus position, which converge into single crested lateral moraines on either side and extend up-valley towards the base of Thompson Peak

(Figure 5). The terminus of the M-1 moraine is positioned ~460 meters (lateral distance) down-valley from M-2. The east and west limbs of M-1 are clearly visible on either side of the lake and the toe of M-1 is positioned below water level in the southern half of modern day Grizzly Lake (Figure 5). M-1 appears to represents the most extensive glacial advance of the Grizzly Valley Glacier contained within the cirque. On the west side of the cirque, above Grizzly Lake, a composite moraine, consisting of three sub-ridges was mapped.

5.2 Establishing Relative Moraine Ages

Field descriptions of M-1 and M-2, summarized in Table 2, clearly distinguish a relative chronology of former glacier positions within the cirque. M-2 boulders have angular, unweathered surfaces with little to no lichen growth and the moraine is highly unstable with no soil development. The crest of M-2 is very sharp, indicating a relatively short period of time since the glacier retreated away from the proximal side of the moraine (Figure 13). Clast counts on the crest of M-2 reveal only 6% of the moraine

27 surface is covered with soil (mostly composed of rock flour), while boulders account for

94% of the moraine surface (Figure 14). These boulder and moraine characteristics are in accord with descriptions of the (LIA) Matthes moraines in the Sierra Nevada range

(Bowerman and Clark, 2005) and thus indicate that M-2 was deposited relatively recently

(100-200 ybp), during the Little Ice Age.

The M-1 moraine shows many indications of its older depositional age, relative to

M-2. M-1 boulders are predominantly sub-rounded with minor feldspar degradation. The moraine is a relatively stable geomorphic feature and the surface has substantial soil development and tree growth (Figure 15 and Table 3). The crest of M-1 is broad and rounded (Figure 13). Clast counts on the crest reveal 47% of the moraine surface is soil, while 53% is exposed boulders (Figure 16), significantly larger than the soil:clast ratio found on M-2. The soil on M-1 consists of pine needles and silty loam ~20-60 cm thick.

The relative ages of M-1 and M-2, based on boulder and moraine characteristics, are consistent with respective positions of the two moraines in the cirque. M-2 is inset ~460 meters within M-1, clearly indicating M-2 was deposited more recently than M-1. The three ridges of the composite moraine on the west side of the cirque are similar in appearance to M-1, with significant soil development consisting of silty loam and decomposed pine needles, substantial tree growth, broad crests, sub-rounded to rounded boulders, and a structurally stable surface. These moraine and boulder characteristics of the three recessional ridges on the west side of the cirque suggest it is similar in age to M-

5.3 10Be Ages

Thirteen rock samples were taken from the Grizzly Valley cirque, seven of which were processed for 10Be cosmogenic radionuclide surface exposure analysis (Figure 5 and

17). The ages of these samples range from 14035 ±1319 yr to 75 ±19 yr. 10Be ages applied to moraines M-1 and M-2 were derived from moraine boulders located on moraine crests. Assuming no post-depositional boulder movement, it is assumed these ages reflect the time since ice receded from the moraine, subsequently ending any additional sedimentation or boulder deposition on the moraine.

Five samples were collected from boulders on the M-1 moraine, two of which were processed for cosmogenic dating. The 10Be ages for these two boulders are 14035 ±1319 and 2634 ±252 yrs (Figure 18). Ages from the M-2 moraine range from 6142 ±577 to 75

±19 yrs (Figure 18). The M-2 cosmogenic ages have an arithmetic mean of 1984 yrs.

However, because M-2 is a composite moraine, samples from the inner and outer crests should be quantified separately. The ages on the inner M-2 crest are 75 ±19 yrs and 198

±29, with an arithmetic mean of 137 yrs. The ages from the outer M-2 moraine are 6142

±577, 2491 ±236, and 1016 ±107 yrs (Figure 18). The mean age for the samples from the outer M-2 crest is 3216 years. The 10Be age of the glacially striated bedrock sampled within M-2 is 5206 ±494 yrs (Figure 18). The bedrock sample collected on the exposed granite between the Grizzly Lake and Grizzly Falls has a 10Be age of 10845 ±1026 yrs.

5.4 Dendrochronology

The oldest tree growing on M-1 is 620 years old. With the addition of the ecesis interval (32-96 yrs), the minimum age of M-1 is 716-652 years old (Figure 19). The oldest tree growing on M-2 is 71 years old. After the addition of the ecesis interval, the

M-2 moraine age is 167-103 years (Figure 19). Assuming that the ecesis time for growing on M-1 and M-2 are the same, a 549 year interval existed between the deposition of M-1 and M-2. The mean age of sampled trees on M-1 and M-2 are 330 and 48 years, respectively. The graph of standardized tree-ring chronologies (index values) for the east and west limbs of M-1 show exceptionally wet and dry episodes over the last ~300 years

(Figure 20). The tree-ring width records from M-1 trees, based on the COFECHA analysis, show increased precipitation during several periods over the last ~200 years including 1807-1818, 1827-1841, 1853-1865, and 1875-1885 (Figure 21).

5.5 Equilibrium Line Altitudes

To establish the modern equilibrium line altitude of the Grizzly Valley Glacier with the accumulation area ratio, we must assume the glacier is in steady state equilibrium.

The calculated ELA for the modern glacier is 2520 ± 11 m (Figure 22). This altitude was

2 established based on a total glacier area of 35004 m and an accumulation area of 22780

2 m (AAR = 0.65 ± .05:1). The ELA of the M-2 glacier extent is 2453 ± 20 m, with a total

2 2 glacier area of 166177 m and an accumulation area of 108100 m (AAR = 0.65 ± .05:1)

(Figure 22). The reconstructed ELA for the M-1 glacial position is 2360 ± 35 m,

2 corresponding to a total glacier area of 744593 m and an accumulation area of 483948

2 m (AAR = 0.65 ± .05:1). The change in ELA, known as the ELA depression, is the difference in altitude between reconstructed glacier equilibrium lines and is the basis for determining the magnitude of climate change necessary to reach variable positions of glaciers in steady state equilibrium. The ELA depression between the modern glacier

(2520 m) and M-2 (2453 m) is calculated to be ~67 m. The ELA depression between the modern glacier (2520 m) and M-1 (2360 m) is ~160 m (Figure 22).

5.6 Climate Change Estimates

Climate records suggest glacier mass variability is influenced primarily by summer temperature and winter precipitation (Porter, 1977). In order sustain the Grizzly Valley

Glacier at the M-2 extent, mean summer temperature would need to decrease by ~0.4°C and winter snow accumulation would need to increase by ~44 cm (Table 4). To sustain the glacier at the M-1 position, mean summer temperature would need to decrease by

~0.9°C and winter snow accumulation would need to increase by ~95 cm (Table 4 and

Figure 23).

Chapter 6: Discussion

6.1 Ages of Moraines and Striated Bedrock

6.1.1 10Be Ages

Five of the 10Be ages sampled from boulders on M-1 (14035 ±1319 and 2634 ±252 yrs) and M-2 (6142 ±577, 2491 ±236, and 1016 ±107 yrs) have exceptionally old ages, compared to the relative dating and dendrochronology results, which are likely due to inheritance of 10Be. Dating moraine boulders based on cosmogenic radionuclide accumulation assumes that the glacial advance responsible for transporting the boulders produced sufficient glacial erosion of bedrock to remove the cosmogenic radionuclides

(10Be in our investigation) accumulated during previous exposure at the surface. If sufficient glacial erosion does not occur when a glacier is advancing or in equilibrium to remove all previously accumulated 10Be, or if boulders have had previous exposure prior to their deposition within the moraine, then inheritance will be an issue for dating moraine boulders and striated bedrock. Rocks previously exposed to cosmogenic

31 radiation on cirque walls and subsequently transported via rockfall, as boulders, onto the glacier or beyond its margin, present another potential source of 10Be inheritance. Also, if the sampled boulder was not eroded from the bedrock during the glacial advance that deposited the moraine, which would be the case when a glacier re-works boulders from older moraines, inheritance will cause the boulder age to be older than the age of moraine deposition.

Both of the 10Be ages from sampled boulders on M-1 have exceptionally old ages

(14035 ±1319 and 2634 ±252 yrs), most likely due to inheritance. The two exceptionally old 10Be ages from M-1 could be related to regional glacial events, such as the Last

Glacial Maximum (LGM) and an advance recognized throughout the Sierra Nevada around 3000 ybp (Figure 17) (Bowerman and Clark, 2005; Konrad and Clark, 1998).

None of the 10Be ages sampled in the Grizzly Valley cirque are older than the LGM (~18 ka), meaning the LGM advance was likely the last advance to cause sufficient erosion to remove all previously accumulated 10Be from the cirque bedrock. The 14035 ±1319 year old boulder from M-1 may have been transported during the LGM advance ~18 ka and deposited on the valley floor during the subsequent recession (~15 ka) or was exposed to the surface on the cirque wall after LGM erosion. Therefore, the boulder would have started accumulating 10Be as it sat on the valley floor or after it was exposed on the cirque wall. The subsequent late-Holocene advance would have mobilized the boulder with inherited 10Be, depositing it in the M-1 moraine. It is likely that the 2634 ±252 year old boulder was also deposited on the valley floor during the recession of the documented advance ~3000 ybp in the Sierra Nevada (Bowerman and Clark, 2005; Konrad and Clark,

1998), before being re-transported and deposited in the M-1 moraine. Another possible

32 explanation for the 10Be inheritance is that the boulders (14035 ±1319 and 2634 ±252 yrs) could have fallen from the cirque walls onto the glacier or up-valley of the M-1 maximum extent, after which they were remobilized and incorporated into the M-1 moraine. A third possible explanation for the inheritance is that the side of the boulder from which the sample was taken was the side of cirque floor bedrock that was exposed to cosmogenic nuclide bombardment, prior to being eroded (plucked) and deposited by the glacier into M-1.

10 Be ages from the outer M-2 samples display a relatively high degree of variation, making it difficult to establish any confident depositional ages for the moraine (6142

±577, 2491 ±236, and 1016 ±107 yrs). There are a number of possible explanations for these unreasonably old and inconsistent 10Be ages. The 2491 ±236 yr boulder may have been eroded from the bedrock during an advance equivalent to the 3000 ybp advance in the Sierra Nevada, discussed in the previous paragraph (Figure 17), before being retransported to the outer M-2 position. At the time the glacier advanced to its outer M-2 extent, many boulders were likely already deposited on the bedrock of the cirque floor, beyond the ice margin, carried partially down the pre-existing glacier and deposited on underlying bedrock during subsequent glacial retreat. These boulders were probably transported and re-deposited by the glacier during the advance that formed the outer M-2 crest, consequently incorporating the boulders with significant inheritance into the outer

M-2 moraine.

As expected from the relative positions of the two crests, the ages from the inner crest of M-2 (75 ±19 yrs and 198 ±29) are significantly younger than those on the outer crest (6142 ±577, 2491 ±236, and 1016 ±107 yrs). The advance responsible for

33 depositing the outer M-2 crest likely transported and incorporated all previously deposited boulders into the outer M-2 moraine. A subsequent readvance to the inner M-2 extent would then only incorporate boulders recently deposited onto the bedrock or eroded from under the glacier during the advance to the inner M-2 position. This scenario is the most likely explanation for the younger ages on the inner M-2 crest. The cosmogenic sample collected from the glacially striated bedrock within M-2 should theoretically provide an accurate age of the most recent glacial advance, if >2 m of the cirque floor was eroded when the glacier was present, resulting in the deposition of the inner M-2 moraine (LIA). However, the age is significantly older than M-1 and M-2, suggesting ice volume was not sufficient during the M-1 or M-2 extent to cause enough sub-glacial erosion of the bedrock to remove previously accumulated 10Be. The inherited

10Be accumulated in the sampled bedrock during periods when the glacier receded up-valley of the bedrock, leaving it exposed to cosmogenic radiation. The zone of maximum sub-glacial erosion under a cirque glacier is under the equilibrium line altitude, where ice thickness is greatest and ice flow causes maximum friction against underlying material (Dahl et al., 2003). The sample located on the striated bedrock within M-2 is positioned above the zone of maximum erosion during the M-1 paleo-glacier extent. The steep bedrock grade within M-2 and the relatively small volume of ice at the M-1 and

M-2 extent also likely contributed to the insufficient sub-glacial erosion to remove all bedrock with inherited 10Be. The age of the bedrock (5206 ±494 yrs) does overlap with a documented glacial advance in the Sierra Nevada, below the Powel glacier (Konrad and

Clark, 1998), which could be the equivalent advance responsible for removing enough bedrock to remove all previously accumulated 10Be under the Grizzly Valley Glacier,

34 explaining the inheritance age (Figure 17).

The 10Be age of the striated bedrock north of Grizzly Lake, near the lip of the cirque, indicates the Grizzly Valley Glacier advanced beyond the lip of the cirque around

10845 ±1026 ybp. Clearly defined striations in the bedrock at the sample location suggest minimal erosion has taken place subsequent to sub-glacial erosion. This age (~11 ka) corresponds to unpublished 10Be moraine ages sampled near the head of the nearby Sugar

Pine Valley, by Richard V. Heermance, which were dated around 11000 ±1100 ybp. The ages from the Sugar Pine Valley and our ages from the Grizzly Valley suggest there was a significant glacial advance around 11 ka. During this advance, the Grizzly Valley

Glacier must have been thick enough to fill the cirque and remove all bedrock with previously accumulated 10Be from the cirque floor.

6.1.2 Dendrochronology

The depositional ages of M-1 (716-652 yrs) and M-2 (167-103) determined from the oldest living tree growing on the moraine with the addition of the ecesis interval range assumes the oldest tree on the moraine was sampled. While no older trees that were dead during sampling were recognized, the possibility remains that older trees could have fallen and/or been partially decomposed (particularly on M-1), making the minimum moraine age estimates too young. However, if the contemporary trees are a second growth population, the first, now dead population of trees, would have been recognizable in the field, which they were not. Furthermore, the degree of soil development on M-1 agrees with the depositional age determined through dendrochronology (716-652 yrs). Although the M-1 moraine appears on the surface to be composed primarily of soil (Figure 13), the soil actually only covers the surface of the

35 moraine. The trees growing on M-1 have been dropping their leaves (needles) onto the surface of the moraine for around 600 years, based on the tree ages. These needles have accumulated and decomposed on the moraine surface, forming a ~6 inch layer of detritus and soil (Figure 15), beneath which are moraine boulders.

The ages applied to the moraines from tree-ring dating the oldest tree have some potential uncertainty related to absent rings, meaning missing years in the tree-ring record resulting in an underestimate of moraine age. However, absent rings in the oldest tree would not significantly affect the moraine age and would be outweighed by the range of ecesis interval added to the tree age. Rather than applying a distinct ecesis time to the tree age, which is a highly variable value, we applied an ecesis range to the tree ages taking into consideration previous dendrochronology investigations and the rather harsh environment of the M-1 and M-2 moraines. The tree-ring age determinations were made assuming the trees had no absent rings, which can occur due to periods of particularly harsh environmental conditions, such as lack of precipitation. If the sampled trees failed to produce rings some years, the tree age and related minimum moraine age would be underestimated.

6.2 Glacier Dynamics and Growth

6.2.1 Glacier Response to Climate

Small cirque glaciers, such as the Grizzly Valley Glacier, are particularly sensitive to climate fluctuations and are capable of responding to these changes in less than a decade (Lillquist and Walker, 2006; Porter, 1986). Small changes in winter precipitation and/or summer temperature have a greater influence on the mass balance of small mass, mid-latitude glaciers than to high latitude valley glaciers, which have prolonged response

36 times (Porter, 1986). Due to the fast response time of small glaciers to climate forcings, composite moraines comprising recessional sub-ridges, with short (decadal scale) depositional time differences are often deposited beyond the ice margin of these glaciers

(Bowerman and Clark, 2005; Schaefer et al., 2009; Schimmelpfennig et al., 2012).

Therefore, the two ridges comprised in the M-2 composite moraine beyond the Grizzly

Valley Glacier could be very close in depositional age (<20 years) and reflect only minor variations of the ELA and climate.

6.2.2 Equilibrium Line Altitude Fluctuations

The accuracy of calculating the modern ELA and paleo-ELAs of M-1 and M-2 is essential in this study, because the estimated climate change during the deposition of these two moraines is largely based on the ELA calculations. The accumulation area ratio method chosen to calculate the ELAs seems to be the best fit for the Grizzly Valley

Glacier, due to the irregular shape and high grade of steepness between Thompson Peak and Grizzly Lake. Previous investigations have focused on the reliability and appropriate applications of different methods for calculating ELA (Meierding, 1982; Osipov, 2004).

Meierding (1982) carried out a study based on moraines in the Front Range in Colorado comparing the reliability of different methods for determining paleo-ELAs, where he evaluated the following methods: maximum elevation of lateral moraines, toe-to-headwall ratio, median elevation of glaciers, cirque-floor, and accumulation area ratio (AAR). The study concluded that the AAR method provided the most reliable ELA results among the various methods. Osipov (2004) also found that the AAR method provides the most satisfactory results on a physical basis, including small glaciers, because it accounts for glacier surface area and absolute elevations. However,

(Bowerman and Clark, 2005) used the maximum elevation of lateral moraines (MELM) method to determine the ELA for small, late Holocene cirque glaciers in the central

Sierra Nevada. The MELM method assumes that due to the nature of glacier flow, lateral moraines do not form above the ELA and post-depositional erosion is negligible

(Bowerman and Clark, 2005; Dahl et al., 2003). In the case of the Grizzly Valley cirque, the highest extent of the east limbs of M-1 and M-2 project onto very steep bedrock, where the moraine is not well defined. The unclear boundary of the uppermost moraine extent on M-1 and M-2 make the MELM method a poor option for our study area.

The uncertainties for the calculated ELAs are based on the error incorporated in the accumulation area ratio formula (AAR= 0.65 +/- .05:1). Because the Grizzly Valley

Glacier is characterized as a small cirque glacier, highly sheltered within cirque walls, the

ELA error could be larger than the assumed accumulation area ratio of 0.65 +/-.05:1. It is possible that the calculated ELAs for the modern and paleo-glaciers are not located precisely where the glacier mass balance equals zero. However, the change in ELA relative to the modern ELA is most important for estimating past climate conditions, not the absolute ELA.

6.3 Climate Record and Glacier Implications

6.3.1 Climate Signal from Dendrochronology

Dendrochronology provides insight into climate conditions during tree growth, through tree-ring widths. The cores sampled from trees growing on the M-1 moraine provide tree-rings extending back 610 years, with the last ~300 years represented in enough cores to be considered an actual climate record proxy, suitable for making climate interpretations. Generally, wide annual tree-rings in our study area correlate with periods

38 of increased precipitation, evident from the concurrent periods of high river discharge in the Trinity River, which can be extrapolated back in time to indicate likely periods of positive glacier mass balance (Figure 21). Consequently, a detailed record of climate conditions during the deposition of M-2 is recorded in the annual tree-ring widths. When the tree ring measurements from the M-1 trees were statistically analyzed with the

COFECHA program to construct a master chronology, the interseries correlation among the cores was fairly low (0.276-east limb and 0.285-west limb) (Table 5). However, this low interseries correlation value, indicative of variable annual tree growth rate between trees, is not surprising considering the trees are growing on dynamic, unstable landforms.

The M-1 and M-2 moraines have highly variable steepness, soil development, soil composition, and subsurface permeability. The heterogeneous growing conditions on the moraines create an environment that is likely to create variable tree-growth rates, even when subjected to a uniform climate regime. The resulting variability in growth rate is a likely cause for the low inter-correlation value among the core subsets analyzed in

COFECHA (Table 5). However, when the index values (reflecting the deviation from mean tree-ring width among the subsets of cores) from the east and west limbs of M-1 are plotted against one another, the periods of thin and wide rings roughly correlate between the two limbs of the moraine. The correlation coefficient between the index values from the east and west limbs is 0.126, which is low, but the anomalously wet and dry periods between the two moraine limbs generally coincide with one another. The tree-ring width record from M-1 also correlates with river discharge data from the nearby Trinity River, available since 1909 (USGS – Online National Water Information System). The plot of the tree-ring width record shows the peaks and troughs correlate through time with

39 periods of high and low river discharge (Figure 21). Therefore, we can assume that this relationship is consistent over the entire history of the tree-ring record, meaning years with relatively wide rings reflect periods of increased rainfall and years with relatively narrow rings reflect dry episodes.

Master chronologies of tree ring widths were constructed using data from nearby

Mt. Eddy (Graumlich, 1981) and Fryday Ridge (Briffa and Schweingruber, 1983), located west and east of the Grizzly Valley, respectively (Figure 24). The Mt. Eddy chronology includes 24 cores and provides a reasonable series intercorrelation (0.485) among tree cores to confidently make interpretations of past climate conditions based on tree ring widths (Table 5). The Fryday Ridge tree-ring width data yields an intercorrelation value of 0.693, indicating another acceptable dataset from which to make climate interpretations (Table 5). The periods of wide and narrow rings from M-1 tree cores roughly correlate with tree-ring width data from Mt. Eddy and Fryday Ridge

(Figure 21). Therefore, even though our tree cores yield a relatively low interseries correlation value, our record appears to be an accurate proxy for the local precipitation record of the Grizzly Valley.

From around 180-135 ybp, the standardized chronologies from the east and west limbs of M-1 go out of sync with one another (Figure 20c). When the tree ring widths from Grizzly Valley are corrected by adding 11 years throughout this period, which would account for possible absent rings, the tree ring widths continue show a much stronger correlation with one another (raising the correlation coefficient from 0.126 to

0.174), as well as with the chronologies from Mt. Eddy and Fryday Ridge, as seen in

Figure 21. The 11 missing years were most likely interspersed throughout the chronology

40 from 180-135 ybp, not 11 consecutive missing rings. If the trees on M-1 had failed to produce annual rings during this period, due to environmental factors, the trees from M-1 would in fact be 11 years older than the age determined by counting the tree-rings (716-

652 ybp). Another possible cause of absent rings in the M-1 tree record, and therefore an underestimation of moraine age by 11 years, is tree disturbance by avalanche activity.

Avalanches are not unlikely considering the steep slopes of the cirque above the moraines. It is possible that destructive avalanche activity could have destroyed trees growing on M-2, previous to the trees sampled in this study, meaning the moraine ages we determined are too young.

6.3.2 Climate Chronology

The master chronology of tree-ring widths from skeleton plots is fairly inconclusive and does not show definitive trends in wet-dry periods. Alternatively, the master chronologies of the trees growing on the M-1 east and west limbs, created with

COFECHA, show a general correspondence with one another and correspond with water discharge of the nearby Trinity River, providing a precipitation record with annual resolve (Figure 21). Tree-ring records from Mt. Eddy (Graumlich, 1981), Fryday Ridge

(Briffa and Schweingruber, 1983), and our investigation in the Grizzly Valley cirque, provide a high-resolution proxy for the local precipitation record. By plotting a 7 year running mean of the three tree-ring master chronologies (Grizzly Valley, Mt. Eddy,

Fryday Ridge), significant positive deviations from the mean (wet episodes) occur from

1805-1818, 1825-1841, 1850-1868, and 1875-1885 (Figure 21). This detailed record of precipitation patterns provides precise time constraints on the climate conditions responsible for mass fluctuations of the Grizzly Valley Glacier. These tree-ring records

41 indicate that the periods of increased precipitation in the Grizzly Valley cirque over the last ~250 years were concurrent with Northern Hemisphere and global temperature depressions, evident in tree-rings, ice cores, historical data, glacier lengths, and borehole data (Moberg et al., 2005; Crowley and Lowrey, 2000; Leclercq and Oerlemans, 2012;

Huang, 2004) (Figure 25). These cold, wet intervals were likely responsible for glacial extension and periods of positive glacier mass balance.

Graumlich (1987) reconstructed annual precipitation patterns for three regions within the Pacific Northwest, based on 41 tree ring chronologies (Figure 26 and 27). She found that in northern California and southern Oregon, including our study site, wet climate conditions occurred from 1850 to1895 with maximum precipitation from

1850-1870 (Figure 28). Graumlich (1987) also concluded that long term precipitation variations in northern California and southern Oregon were not synchronized with the rest of Oregon and Washington, which supports the findings in this thesis that the chronology of late-Holocene glacier fluctuations in the Klamath Mountains is out of synch with ranges to the north and south. Keen (1937) used 265 stump cross sections in southeast Oregon to construct a precipitation record of the region based on tree-ring widths. The results show tree-growth rates were anomalously high between 1855 and

1870, indicating higher than average precipitation during this period. The timing of this relatively wet episode correlates with the mid-nineteenth century precipitation increase documented by Graumlich (1987) from 1850-1870 (Figure 28). This ~20 year period of increased precipitation is also represented in the tree-ring records from our investigation in the Grizzly Valley cirque, Mt. Eddy (Graumlich, 1981), and Fryday Ridge (Briffa and

Schweingruber, 1983). Therefore, ample evidence suggests our study area, along with

42 southern Oregon and the rest of northern California, experienced a cool, wet climate between 1850 and 1870.

6.3.3 Climate Change Estimates

Weather stations chosen to calculate the local precipitation lapse rate were chosen based on proximity to the study area, altitude, and similar longitudinal position (Table 6 and Figures 12 and 29). A power regression curve was chosen to fit to the precipitation data points because it appears to provide the most realistic precipitation values when extrapolated to higher elevations. A power regression curve was also chosen by

Bowerman and Clark (2005) to determine a local precipitation lapse rate in the central

2 Sierra Nevada, California. In our study, a power regression curve fits the data with an R value of 0.9716 (Figure 29). Temperature data from the weather stations chosen to determine a local temperature lapse rate clearly display a linear trend. The linear

2 regression line fit to the data has an R value of 0.9754. Local temperature lapse rate calculated for the Sierra Nevada (Bowerman and Clark, 2005) and the Three Sisters

Volcanoes, Oregon (Marcott, 2005) is similar to that calculated for in our study area

2 (Figure 30). Based on our R values (0.9754 for temp. and 0.9716 for precip.) and the comparison to other calculated lapse rates in western North America (Figures 28)

(Bowerman and Clark, 2005; Marcott, 2005), the precipitation and temperatures lapse rates determined in our investigation seem highly reasonable.

As discussed, the calculated ELAs in this study have some inherent uncertainty and these uncertainties were projected into calculating summer temperature and winter precipitation at the various ELAs (Table 6). Climate conditions at the modern Grizzly

Valley Glacier ELA plot to the right of modern glacier climate envelope developed by

Leonard (1989), indicating the modern glacier exists under anomalous summer temperature and winter precipitation conditions (Figure 23). The plot of the Grizzly

Valley Glacier on this chart highlights that the Grizzly Valley cirque, as well as the

Klamath Mountain region, has a warm and exceptionally wet climate, relative to most glacier supporting regions. The magnitude of change necessary for climate conditions at the reconstructed M-1 and M-2 ELAs to reach the conditions at the lower global ELA climate envelope boundary was not calculated, because a contemporary glacier is available for climate-elevation comparison. Determining the necessary winter accumulation increase and mean summer temperature decrease to attain the global ELA conditions would overestimate the severity of climate during the M-1 and M-2 glacier extent (Gillespie and Zehfuss, 2004). Burbank (1986) claims that a temperature would need to decrease approximately 0.5-1.2°C from current temperature for a glacier to advance to the typical LIA maxima. He also states that if these temperature depressions were concurrent with increased winter accumulation, less of a temperature decrease would be necessary for glaciers to reach common LIA extents. In our investigation, a temperature depression of ~0.4°C (with a winter accumulation increase of ~44 cm)

(Table 4) was calculated for the M-2 glacier position, which is on the lower end of

Porter’s (1986) proposed temperature depression range, however our tree-ring records show our study area did in fact receive increased precipitation during significant temperature depressions around the time of M-2 maximum extent. Our results show a temperature depression of ~0.9°C and a 95 cm increase in winter precipitation (Table 4) is needed for the glacier to reach the M-1 extent, which seems reasonable considering

0.9°C falls near the middle of the temperature depression range proposed by Porter

(1986). The significant increase in winter precipitation required to grow the glacier to the

M-1 position, relative to the depression in summer temperature needed, suggests that precipitation is the primary contributing factor to the mass balance of the Grizzly Valley

Glacier. These are the first quantitative climatic constraints applied directly to glacial events in the Klamath Mountains; therefore a comparison within this particular region is not possible. However, glacier-climate investigations have been carried out in the

Cascade and Sierra Nevada ranges (Bowerman and Clark, 2005; Marcott, 2005), which are useful for comparison.

Bowerman and Clark (2005) found that in the central Sierra Nevada, mean summer temperature would have been ~0.2-2.0°C lower than today, and winter accumulation would need to increase by ~3-26 cm (SWE) for modern glaciers to reach the Matthes extent (equivalent to our M-2 extent). Marcott (2005) found that for current climate conditions at the modern glaciers of the Three Sisters Volcanoes, in the Cascades, to reach those during the LIA, mean summer temperature would need to decrease by

0.2-1.0°C combined with a winter accumulation increase of 10-60 cm (SWE). The large range of temperature variations in the mentioned studies is because their calculations were made using multiple glaciers and therefore, multiple values for ELA depression.

The difference in LIA temperature calculated in our investigation (0.4°C) is within the range determined both in the central Sierras and the southern Cascades (Bowerman and

Clark, 2005; Marcott, 2005). Alternatively, the estimated change in LIA precipitation for the Grizzly Valley Glacier (~44 cm) is higher than the range determined for the Sierras

(~3-26 cm), while it is within the range for the southern Cascades (10-60 cm). The relatively high change in precipitation agrees with our explanation for the anomalously

45 low altitude existence of the Grizzly Valley Glacier (anomalously high local precipitation).

6.3.4 Timing of Moraine Deposition

6.3.4.1 M-1

The 10Be cosmogenic ages sampled on M-1 are highly variable, but all post-date the

LGM (~18 ka). On the M-1 moraine, the 10Be boulder ages do not provide much insight on the absolute depositional age of M-1, unlike the tree-ring ages. Because tree ages can be determined with high precision, by counting annual tree rings, dendrochronology provides the most satisfactory moraine age estimates for this investigation. The primary uncertainty in assigning a landform an age based on a tree age is the ecesis interval, or time from moraine stabilization to seed germination. Here, we apply an ecesis interval range of 32-92 years to tree ring counts. Tree ages, with the addition of the ecesis interval, constrain the minimum depositional age of M-1 to between 716-652 ybp. This age range can be better constrained by observing local precipitation and temperature trends around the depositional age determined by tree ages (716-652 ybp). Northern

Hemisphere temperature proxies indicate significant temperature depressions in the

Northern Hemisphere, potentially associated with the M-1 glacier advance, from around

760-710 ybp (1250-1300 C.E.) and 690-647 ybp (1320-1363 C.E.) (Figure 25). These temperature records were constructed using tree-ring and ice-core data (Crowley and

Lowrey, 2000; Moberg et al., 2005). Considering the age range of M-1 deposition based on the dendrochronology investigation (716-652 ybp), the temperature record suggests that the cool period from 760-710 ybp (1250-1290 C.E.) is the most likely climate episode responsible for shifting the Grizzly Valley Glacier into positive mass balance,

46 causing the glacier to advance to the M-1 position during this period. The subsequent warming trend from around 710-700 ybp (Figure 25) would have resulted in glacial retreat and the deposition of M-1, making the preferred depositional age of this moraine around 690 ybp. Although the cool period, based on the ice-core and tree-ring data, from

690-647 ybp overlaps the M-1 age determined from dendrochronology more than the

760-710 ybp cool interval, the timing of termination of a cool interval is most important when determining the age of moraine deposition.

The composite moraine to the west of the Grizzly Valley Glacier consists of a terminal moraine above Grizzly Lake with two inset moraines (Figure 5), which have similar surficial characteristics to the M-1 moraine (soil development, tree growth, stable landform, broad crest). Therefore, the outer crest of the composite moraine (Mw-1 in

Figure 5) is most likely the same age as M-1. The glacier lobe that deposited the M-1 moraine likely experienced the same mass balance fluctuations as the lobe on the west side of the cirque, and therefore should theoretically have deposited the two inset moraines (older than the M-2 moraine) within the M-1 position. However, the cirque floor is much steeper within the M-1 moraine, relative to the slope within the western composite moraine, making moraine preservation within M-1 highly unlikely.

An alternative explanation for the lack of moraines between M-1 and M-2 is that the glacier lobe, which deposited M-1 behaved differently than the lobe that deposited the composite moraine on the west side of the cirque. The glacial fluctuations between the

M-1 and M-2 moraines may have not decreased in magnitude successively, as it appears the western lobe did, evident from the two inset ridges. If the three ridges of the composite moraine on the west side of the cirque are recessional ridges, it is possible the

M-1 glacier lobe did not recede in a similar, episodic fashion. The glacier lobe that deposited M-1 would be expected to behave somewhat differently than the lobe responsible for depositing the moraines west of the modern glacier (Mw composite), because the M-1 lobe received the most topographic shielding from sunlight due to

Thompson Peak, which is one reason the Grizzly Valley Glacier exists today. When the glacier extended to the M-1 position, it would have transported and therefore removed all depositional traces of the previous glacier fluctuations, if they existed. If the glacier fluctuated between the M-1 and M-2 position, prior to the deposition of M-1, tree growth would have been inhibited to the modern M-1 moraine.

6.3.4.2 M-2

M-2 is a composite moraine consisting of two closely spaced crests (<22m). The inner and outer crests of M-2 reflect at least two fluctuations of glacier mass balance during the Little Ice Age. The detailed climate record provided by tree rings allows us to correlate distinct climate episodes with glacier fluctuations, thereby constraining the timing of glacial advance and retreat likely for depositing the two M-2 moraines. There are multiple scenarios in which the glacier could have responded to these climate oscillations, resulting in the deposition of the M-2 inner and outer crests (primarily readvancing or pausing during ice retreat). However, by comparing the available M-2 moraine age range with the climate history in the tree-ring record, a most probable

(preferred) depositional history for M-2 can be established.

Maximum tree age (71 yrs) and ecesis interval (32-96 yrs) indicate the outer M-2 crest was deposited between 1845 and 1909 C.E. This age range provides maximum and minimum depositional ages of the outer M-2 moraine, based on the range of ecesis

48 interval, but this wide moraine age range is not sufficiently precise, considering the decadal scale of the LIA. Although the ecesis interval for M-2 was likely longer than the mean ecesis of previous investigations (32 yrs), due to the harsh environmental conditions and resistant granite soil source, it was probably not as prolonged as the longest documented ecesis in North America (96 yrs). Here, we generally consider prolonged intervals of high precipitation favorable for glacier growth or steady state equilibrium (net mass balance ≥0), and intervals of less than average precipitation likely periods of glacier retreat (negative net mass balance). This relationship of precipitation and glacier net mass balance provides the basis from which we determine the timing of glacier fluctuations associated with the deposition of M-2.

A period of cold, wet climate conditions, from ~1850-1870 C.E., is well documented through tree-ring studies in the study region, which coincides with the age range of M-2 determined from our dendrochronology study (1845-1909 C.E.) (Figure

25). This ~20 year period of increased precipitation and cool temperatures represents the most likely climate episode responsible for advancing the glacier (or pausing recession) to its outer M-2 extent. This two decade period displays the greatest positive deviation from the mean precipitation within the age range of M-2 and is the most prolonged period of greater than normal precipitation within the age range. Other regional tree-ring records

(Graumlich, 1987; Keen, 1937) characterize this period 162-142 ybp (~1850-1870 C.E.) as a wet episode, offering additional support for the glacier having a positive mass balance and advancing to the M-2 maximum extent. Subsequent retreat, prior to inner

M-2 deposition, likely occurred within the next 5-15 years, as indicated by the local tree-ring record, which shows less than average precipitation. The retreat would have

49 caused the glacier ice to abandon the surface of outer M-2, bringing an end to all glacial deposition. Therefore, the preferred age of the outer M-2 moraine is around 140 ybp

(1872 C.E.). However, it is noteworthy that the modern Grizzly Valley Glacier mass balance has been sustained for at least 100 years, based on air photos, and there are no moraines at the ice margin. Therefore, it is possible that the glacier was at static equilibrium at the M-2 position for over 200 years before receding from the outer M-2 position, especially considering the volume of boulders deposited in the outer and inner

M-2 moraines (assuming the glacier at the M-2 extent eroded all the boulders from the bedrock rather than mobilizing them from previous deposits).

Another period of increased precipitation occurred from 1875-1885 C.E., which coincided with a period of northern hemisphere temperature depression (Crowley and

Lowrey, 2000; Huang, 2004). This cool, wet interval likely resulted in a transfer of glacier net mass balance from negative (receding) to positive (advancing) or zero (steady state equilibrium). If the increased precipitation caused the glacier to enter steady state equilibrium, essentially pausing glacier retreat, the inner M-2 crest would be characterized as a recessional moraine. Alternatively, the glacier may have retreated up-valley, beyond the modern day M-2 inner crest position, before readvancing to its M-2 inner position. In either case, the glacier likely retreated rather continuously once the recession initiated from the M-2 inner extent (post-1885 C.E.), when winter precipitation would have been insufficient to balance out summer ablation. The absence of glacial till between inner M-2 and the modern glacier supports the idea of an uninterrupted retreat after 1885 C.E. Therefore, the preferred age for the inner M-2 moraine is ~127 yrs (1885

C.E.).

6.4 Regional Significance

6.4.1 Regional Perspective

The research detailed in this thesis suggests that the climate patterns in the Klamath

Mountains during the late-Holocene were asynchronous with the rest of western North

America. The existence of the modern Grizzly Valley Glacier at anomalously low altitude alone, is indicative of the unique climate conditions this area experiences, considering the glacier exists 610 meters lower than any contemporary glaciers in the nearby Sierra Nevada range (Guyton, 1998). The dendrochronology and cosmogenic results from our investigation indicate that the unique climate regime in our study area resulted in late-Holocene glacier fluctuations that were different in timing and magnitude than the fluctuations in the Sierra Nevada and Cascades ranges. The unique climate of the

Grizzly Valley region, which currently sustains (and possibly adds) glacier mass, is also evident at nearby Mt. Shasta. The Whitney and Hotlum Glaciers (elevations of 4200-

3000m and 4100-3200m, respectively), on nearby Mt. Shasta (~50km away), have advanced since the mid-twentieth century in the midst of a regional warming trend of several degrees Celsius (Howat et al., 2007). The glacier growth on Mt. Shasta supports our hypothesis that glacier mass balance in this region is primarily influenced by winter precipitation, rather than summer temperature.

Tree-ring investigations from Mt. Eddy (Graumlich, 1981), Fryday Ridge (Briffa and Schweingruber, 1983), and our investigation in the Grizzly Valley show cold, wet periods from 1805-1818 and 1825-1841 C.E., which overlap with documented LIA glacial maxima in the Sierra Nevada (Matthes advance) and Cascade ranges. It is likely the Grizzly Valley Glacier had a positive mass balance during these wet episodes,

51 causing the glacier to advance. However, while the LIA glacial maxima in the Sierra

Nevada and Cascade ranges appear to end around 1850, initiating glacial retreat, the climate record of the Grizzly Valley suggests it sustained its volume until ~1885.

Therefore, our results indicate the cool, wet climate conditions, possibly linked with those that caused the glacial advance in the Sierra Nevada (Matthes advance) and the

Cascades between around 1783-1850 C.E., persisted in the Klamath Mountains at higher elevations until the late nineteenth century (Figure 25). These results suggest large-scale, regional climate regimes can have a variable effect on areas over relatively short lateral distances. Meier et al. (1980) also documented variable climate patterns during the

Holocene in western North America. They found that glacier mass balances in southern

Alaska were inversely related to those in British Columbia and the southern Cascades

(Meier et al., 1980).

The M-1 and M-2 moraines highlight two significant anomalies in the Grizzly

Valley Glacier’s historical hypsometry, including the glacial maximum around 690 ybp and a sustained Little Ice Age glacier that did not diminish significantly in size until after

130 ybp. Both of these are abnormal when compared with glacier mass fluctuations in rest of the western United States. Because moraine ages indicate the time of glacier retreat, it is not possible to definitively say when or how many advances occurred between ~600-200 ybp. As previously discussed, the existence of the Grizzly Valley

Glacier between the elevation of 2460 and 2560 meters is also anomalous in comparison with the altitudes of modern glaciers in the Sierra Nevada and Cascade ranges.

The ~690 ybp glacial advance in the Klamath Mountains, correlates with a well- documented advance in the European Alps, Iceland and Scandinavia (Davis et al., 2009).

Like the Grizzly Valley Glacier, glaciers in the European Alps, Scandinavia and Iceland reached a glacial maximum around 700-600 ybp, which is not apparent in the Sierra

Nevada and Cascade Ranges. The chronology of the late-Holocene Grizzly Valley

Glacier fluctuations seems to be more similar to glacier variations in the European Alps,

Iceland and Scandinavia rather than the proximal ranges in western North America.

Therefore, the timing of late-Holocene glacial maxima in the Klamath Mountains is locally unusual, but relatively normal in hemispheric context. Not only does the documented ~700 ybp advance of hemispheric extent offer some support to our M-1 age determination, it also supports a larger scale LIA forcing mechanism(s), such as solar variability, changes in summer insolation, or deep water circulation variations, rather than a more regional forcing mechanism. More high-resolution glacier-climate reconstructions are needed to better understand the cause of these natural, but recent climate fluctuations.

Although climate interpretations cannot be confidently made from one cosmogenic exposure age, it is noteworthy that this age (10845 ±1026 yrs) overlaps with documented

Younger Dryas advances in western North America: the 11.5+/-0.3 ka in the Canadian

Rockies, Alberta (Reasoner et al., 1994), 11.3 ka the Enchantment Lakes region,

Washington (Bilderback, 2004), and 11.05 ka near Lakedale in the Northern Cascades,

Washington (Porter, 1976). The age also overlaps documented post-Younger Dryas glacial advances in western North America, including a 10.4 ka advance on Mount

Rainier, Washington (Heine, 1997), and a 10.2 ka advance in the Wallowa Mountains,

Oregon (Licciardi et al., 2004). Although advances during the Younger Dryas

Chronozone are evident in the Rocky Mountains and Northern Cascades, no significant glacial advances appear to occur during the Younger Dryas in the Sierra Nevada

(Gillespie and Zehfuss, 2004). If Younger Dryas advances did occur in the Sierra

Nevada, they were less extensive than the Matthes (LIA) advance, which is certainly not the case in our field area, based on the 10.8 ka cosmogenic age. The apparent inconsistency of Younger Dryas advance along western North America agrees with our speculation from the late-Holocene results in this study, that local response to regionally widespread or global climatic events is variable.

6.4.2 Correlation With Regional Cycles

Climate variations in the Pacific Northwest are intimately linked with atmosphere-ocean interactions (Lillquist and Walker, 2006). The Pacific Decadal

Oscillation (PDO) is a well-known example of these atmosphere-ocean circulation interactions, which is a function of sea-surface temperature, and is characterized by alternating warm, dry – cool, wet cycles with a periodicity of 20-30 years (Lillquist and

Walker, 2006; Roe and O'Neal, 2009). Positive phases of the PDO typically coincide with warm, dry winters, while the negative phases are generally cool, wet winters (Koch et al., 2009). During the negative phases of the PDO, the Aleutian Low becomes depressed, redirecting winter storms south of Alaska, which increases precipitation and decreases temperature in the Pacific Northwest (Koch et al., 2009). The PDO patterns appear to correlate with variations in temperature, precipitation, and glacier mass changes on Mt. Hood and Mt. Baker in the Cascades (Lillquist and Walker, 2006; Kovanen, 2003) and on glaciers in Garibaldi Provincial Park, in the southern Coast Mountains (Koch et al., 2009). The precipitation record in our study area, evident from tree-rings, seems to broadly agree with the PDO phases, over the last 300 years, from 1757-1824, 1835-1870, and 1925-1972 (wet episodes corresponding to negative phases and vice versa) (D'Arrigo

54 et al., 2001). Therefore, it appears that the PDO likely had a strong influence on climate patterns in our study area during some periods of the past, but it is likely alternative forcing mechanisms were also largely involved, often dominating the PDO signature.

The climate variations caused by ENSO and PDO are similar, both occurring at a decadal scale and produce changes in sea surface temperature (SST), sea level pressure, and winds, but the PDO has a more widespread affect and is centered on the mid-latitude

North Pacific (Cane, 2005). Bitz et al. (1999) found that mass balance fluctuations in glaciers spanning from Washington to Alaska are significantly affected by atmosphere-ocean circulation anomalies linked with the El Niño Southern Oscillation

(ENSO). Tree-ring widths in our study area also indicate a significant increase in precipitation during the 1982-1983 El Niño event, indicating the ENSO also had an effect on local climate patterns (Figure 21). The extended ENSO record for around the last 300 years shows some periods of strong correlation with tree-ring widths (precipitation) in the

Grizzly Valley cirque, while intervening periods display no such relationship (Figure 31).

Therefore, like the PDO, the ENSO probably plays some role in the climate patterns of our study area, but is one of many forcing mechanisms contributing to the glacier-climate variations during the late-Holocene in the Klamath Mountains.

Conclusions

The Grizzly Valley cirque contains the most detailed record of late-Holocene glacial fluctuations yet to be documented in the Klamath Mountains. The cirque contains scoured bedrock surfaces and well-defined moraines with substantial tree-growth, making this study site an ideal location to investigate the magnitude and chronologic history of

55 variations in climate and glacier mass balance fluctuations during the late-Holocene. The combination of dendrochronology and cosmogenic radionuclide dating methods allowed us to fully utilize the well-preserved moraines, striated bedrock, and tree-growth to quantify the ages of glacier retreat. Because the Grizzly Valley Glacier still exists, it is possible to quantify the temperature and precipitation changes needed to advance the glacier to the position of the well-preserved moraines. Therefore, the combination of dating the age of moraine deposition (M-1 and M-2) and determining the difference in climate conditions during the respective times, makes this research the only documented quantitative glacier-climate investigation in the Klamath Mountains.

10Be cosmogenic dates indicate the Grizzly Valley Glacier filled the cirque around

10.8 ka, possibly coinciding with the Younger Dryas advances in the Cascade Range to the North. Moraine ages, and the climate record from our dendrochronology investigation, indicate three subsequent glacial maxima occurred at ~690, ~150, and ~130 ybp (Figure 32). Local precipitation and temperature lapse rates and ELA reconstructions of the modern glacier and paleo-glaciers suggest that during the ~690 ybp glacial maximum, summer temperature was ~ 0.9°C less and winter precipitation was ~95 cm more than current conditions. Because the ~150 and ~130 ybp maximum glacier extents are so close to one another (< 22 m), the difference in temperature and precipitation patterns during these times is marginal. In comparison to modern conditions, mean summer temperature, around 150 and 130 ybp, was ~0.4°C cooler and winter precipitation was ~44 cm greater. Although Little Ice Age glacial advances occurred in the Sierra Nevada and Cascade ranges, time constraints on the glacial fluctuations in the

Grizzly Valley cirque suggest the timing of these events was variable along western

North America. The ~690 ybp glacier maximum in the Klamath Mountains is not apparent in the Sierra Nevada or Cascade Ranges. Also, glaciers in the Sierra Nevada and

Cascade Ranges receded from their LIA (Recess Peak in the Sierra Nevada) extents ~20-

30 years prior to the glaciers in the Klamath Mountains. The climate in the Klamath

Mountains likely varies from the climate of the Sierra Nevada and Cascade ranges due to the closer proximity to the Pacific Ocean and the associated westerly storm systems. A comparison of documented glacial advances in the Sierra Nevada and Cascade ranges to the chronology of Grizzly Valley Glacier fluctuations suggests regional response to large!scale climate regimes can vary over relatively short lateral distances.

References

Aa, A., 1996, Topographic control of equilibrium-line altitude depression on reconstructed ‘Little Ice Age’ glaciers, Grobareen, western Norway, The Holocene 6, 82-89.

Allen, C.M., and Barnes, C.G., 2006, Ages and some cryptic sources of Mesozoic plutonic rocks in the Klamath Mountains, California and Oregon, in Snoke, A.W. and Barnes, C.G., eds., Geological studies in the Klamath Mountains province, California and Oregon: A volume in honor of William P. Irwin: Boulder, Colorado, Geological Society of America Special Paper 410, 223-245.

Bahr, D., Meier, M., Peckham, S. 1997, The physical basis of glacier volume-area scaling, Journal of Geophysical Research 102(20), 355-362.

Balco, G., Briner, J.P., Finkel, R.C., Schaefer, J.M., 2009, Regional beryllium-10 production rate calibration for late-glacial northeastern North America, Quaternary Geochronology 4, 93-107. Doi:10.1016/j.quageo.2008.09.001.

Balco, G., Stone, J.O.H., Lifton, N.A., Dunai, T.J., 2008, A complete and easily accessible means of calculating surface exposure ages or erosion rates from 10Be and 26Al measurements, Quaternary Geochronology 3, 174-195. Doi:1016/j.quageo.2007.12.001. http://hess.ess.washington.edu/math/al_be_multiple_v22.php

Barclay, J., Wiles, G.C., Calkin, P. 2009, Tree-ring crossdates for a First Millenium AD advance of Tebenkof Glacier, southern Alaska, Quaternary Research 71, 22-26.

Bilderback, E.L., 2004, Timing and Paleoclimatic Significance of Latest Pleistocene and Holocene Cirque Glaciation in the Enchantment Lakes Basin, North Cascades, WA [Masters thesis]: Bellingham, Western Washington University.

Bitz, C.M., Battisti, D.S. 1999, Interannual to Decadal Variability in Climate and the Glacier Mass Balance in Washington, Western Canada, and Alaska, American Meteorological Society 12, 3181-3196.

Bond, G., Kromer, B., Beer, J., Muscheler, R., Evans, M.N., Showers, W., Hoffmann, S., Lotti-Bond, R., Hajdas, R., Bonani, G. 2001, Persistent Solar Influence on North Atlantic Climate During the Holocene, Science 249, 2130-2136.

Bowerman, N.D. and Clark, D.H. 2005, Onset and culmination of neoglaciation in the Sierra Nevada, California, Geological Society of America Abstracts with Programs 37(7), 543.

Briffa, K. and Schweingruber, F.H, 1983, National Climatic Data Center, Fryday Ridge-PSME-ITRDBCA558, http://hurricane.cndc.noaa.gov/pls/paleox/f?p=519:1:::::P1_STUDY_ID:2818.

Burbank, D. 1982, Correlations of Climate, Mass Balances, and Glacial Fluctuations at Mount Rainer, Washington, U.S.A., Since 1850, Arctic and Alpine Research 14, 137- 148.

Burbank, D.W. 1981, A chronology of Late Holocene Glacier Fluctuations on Mount Rainier, Washington. Arctic and Alpine Research 13, 369-386.

Burbank, S.C. 1986. Pattern and Forcing of Northern Hemisphere Glacier Variations during the Last Millenium. Quaternary Research 26, 27-48.

Burke, R.M., 1972, Neoglaciation of Boulder Valley, Mt. Baker, Washington. M.Sc. thesis, Western Washington University, Bellingham. 47p.

Campbell, I.D., Campbell, C., Apps, M.J., Rutter, N.W., Bush, A.1998, Late Holocene ~1500 yr climatic periodicities and their implications, Geology 26, 471-473.

Cane, M.A., 2005, The evolution of El Nino, past and future, Earth and Planetary Science Letters 230, 227-240.

Crowley and Lowrey, 2000, Northern Hemisphere Temperature Reconstruction, Science 289, 270 ftp://ftp.ncdc.noaa.gov/pub/data/paleo/gcmoutput/crowley2000/crowley_lowery2000_n ht.txt.

D’Arrigo, R., Villalba, R., and Wiles, G., 2001, Tree-ring estimates of Pacific decadal climate variability, Climate Dynamics, 18(3-4), 219-224.

D’Arrigo, T., E., Cook, R.W., Allan, R., and Mann, M., 2005, Variability of ENSO over the past six centuries, Geophys. Res. Letters.

Dahl, S., Bakke, J., Oyvind, L., Nesje, A. 2003, Reconstruction of former glacier equilibrium-line altitudes based on proglacial sites: an evaluation of approaches and selection of sites, Quaternary Science Reviews 22, 275-287.

Dansgaard, W., Johnsen, S.J., Clausen, H.B., Dahl-Jensen, D., Gundestrup, N., Hammer, C.U., and Oeschger, H. 1984, North Atlantic oscillations revealed by deep Greenland ice cores, American Geophysical Union Geophysical Monograph 29, 288- 289.

Davis, T., Menounos, B., Osborn, G., 2009, Holocene and latest Pleistocene alpine glacier fluctuations: a global perspective, Quaternary Science Reviews 28, 2021-2033.

Denton, G. and Karlen, W., 1973, Holocene climatic variations – their pattern and possivle cause, Quaternary Research, 3, 155-205.

Easterbrook, D.J. and Burke, R.M., 1971, Neoglaciation on the flanks of Mt. Baker, Washington (abstract). Northwest Science 45, 4.

Fuller, S. R., 1980, Neoglaciation of Avalanche Gorge and the Middle Fork Nooksack River Valley, Mt. Baker, Washington. M.Sc. thesis, Western Washington University, Bellingham 1-68.

Gillespie, A.R. and Zehfuss, P.H. 2004, Glaciations of the Sierra Nevada, California, USA, Quaternary Glaciations – Extent and Chronology, Part II. 51-62..

Graumlich, L.J., 1987, Precipitation Variation in the Pacific Northwest (1675-1975) as Reconstructed from Tree Rings, Annals of the Association of American Geographers 77:1, 19-29.

Graumlich, L.J., 1981, National Climate Data Center/ Paleoclimatology. Mt. Eddy, PIJE, ITRDB CA523, http://hurricane.ncdc.noaa.gov/pls/paleox/f?p=519:1:2929533524396899::::P1_STUD Y_ID:3322.

Grove, J.M., 1988, The Little Ice Age: London, Methuen, 498p.

Guyton, Bill, 1998, Glaciers of California: modern glaciers, ice age glaciers, origin of Yosemite Valley, and a glacier tour in the Sierra Nevada, London, England 197.

Heine, J.T., 1997, Glacier advances at the Pleistocene/Holocene transition near Mount Rainier volcano, Cascade Range, USA: Seattle, University of Washington.

Heusser, C.J., 1957, Variations of Blue, Hoh, and White glaciers during recent centuries, Arctic 10, 139-150.

Howat, I.M., Tulaczyk, S., Rhodes, P., Israel, K., Snyder, M., 2007, A precipitation- dominated, mid-latitude glacier system: Mount Shasta, California, Climate Dynamics 28, 85-98.

Huang, S., 2004, Merging Information from Different Resources for New Insights into Climate Change in the Past and Future, Geophys. Res, Lett. 31, L13205. doi:10.1029/2004GL019781.

Hughes, M.K. and Diaz, H.F. (eds.), 1994, The Medieval Warm Period. Climate Change 26, 109-340.

Jones, Alice G. 1986, Flowers and Trees of the Trinity Alps: A Photographic Guide, Trinity County Historical Society, Weaverville, CA.

Keen, F.P., 1937, Climatic cycles in eastern Oregon indicated by tree rings, Monthly Weather Review 65, 183-88.

Koch, J., Menounos, B., Clague, J. 2009, Glacier change in Garibaldi Provincial Park, southern Coast Mountains, British Columbia, since the Little Ice Age, Global and Planetary Change 66, 161-178.

Kohl, C.P., and Nishiizumi, Kunihiko, 1992, Chemical isolation of quartz for measurement of in-situ -produced cosmogenic nuclides: Geochimica et Cosmochimica Acta 56, 3583-3587.

Konrad, S.K. and Clark, D.H., 1998, Evidence for an early Neoglacial Glacier Advance from Rock Glaciers and Lake Sediments in the Sierra Nevada, California, USA, Arctic and Alpine Research 30, 272-284.

Kovanen, D.J., 2003, Decadal variability in climate and glacier fluctuations on Mt. Baker, Washington, USA, Geofiska Anneler,85A, 43-55.

Kreutz, K.J., Mayewski, P.A., Meeker, L.D., Twickler, M.S., Whitlow, S.I., and Pittalwala, I.I., 1997, Bipolar changes in atmospheric circulation during the Little Ice Age. Science 277, 1294-1296.

Leclercq, P.W. and Oerlemans, J., 2012, Global and hemispheric temperature reconstruction from glacier length fluctuations, Climatic Dynamics 38: 5-6, 1065-1079. Data sourced: http://hurricane.ncdc.noaa.gov/pls/paleox/f?p=519:1:::::P1_STUDY_ID:13544.

Leonard, E.M., 1989, Climatic change in the Colorado Rocky mountains: estimates based on modern climate at late Pleistocene equilibrium lines, Arctic and Alpine Research 21, 245-255.

Lewis, D. and Smith, D., 2004, Dendrochronological Mass Balance Reconstruction, Strathcona Provincial Park, Vancouver Island, British Columbia, Canada. Arctic, Antarctic, and Alpine Research 36, 598-606.

Licciardi, J.M., Clark, P.U., Brook, E.J., Elmore, D., and Sharma, P., 2004, Variable responses of western U.S. glaciers during the last deglaciation: Geology 32, 81- 84.

Licciardi, J.M., Scaefer, J.M., Taggart, J.R., Lund, D.C., 2009, Holocene Glacier Fluctuations in the Peruvian Andes Indicate Northern Climate Linkages, Science 325, 1677-1679.

Lillquist, K. and Walker, K., 2006, Historical Glacier and Climate Fluctuations at Mount Hood, Oregon, Arctic, Antarctic, and Alpine Research 38, 399-412.

Lo, Y.H., Blanco, J.A., Seely, B., Welham, C., Kimmins, J.P., 2010, Relationships between climate and tree racial growth in interior British Columbia, Canada, Forest Ecology and Management 259, 932-942.

Luckman, B.H., 1988, Dating the Moraines and Recession of Athabasca and Dome Glaciers, Alberta, Canada, Arctic and Alpine Research 20, 40-54.

Marcott, S.A., 2005, A Tale of Three Sisters: Reconstructing the Holocene glacial history and paleoclimate record at Three Sisters Volcanoes, Oregon, United States, Portland State University M.S. Ph.D. Thesis 1-93.

Matthews, J., and Briffa, K., 2005, The ‘Little Ice Age’: Re-evaluation of an Evolving Concept, Geografiska Annaler 87, 17-1033.

Mayewski, P.A., Eelco, E., Rohling, J., Stager, C., Karle’n, W., Maasch, K.A., Meeker, L.D., Meyerson, E.A., Gasse, F., can Kreveld, S., Holmgren, K., Lee-Thorp, J., Rosqvist, G., Rack, F., Staubwasser, M., Schneider, R.R., Steig, E.J., 2004, Holocene climate variability, Quaternary Research 62, 243-255.

McCabe, G., Fountain, A., 1995, Relations between Atmospheric Circulation and Mass BAlance of South Cascade Glacier, Washington, U.S.A. Arctic and Alpine Research 27, 226-233.

McCarthy, D.P. and Luckman, B.H., 1993, Estimating Ecesis for Tree-Ring Dating of Moraines: A Comparative Study from the Canadian Cordillera, Arctic and Alpine Research 25, 63-68.

McCarthy, D.P., Luckman, B.H., Kelley, P.E., 1991, Sampling Height-Age Error Correction for Spruce Seedlings in Glacial forefields, Canadian Cordillera, Arctic and Alpine Research 23, 451-455.

Meier, M.F., Mayo, L.R., Trabant, D., and Krimmel, R.M., 1980, Comparison of mass balance and runoff at four glaciers in the United States, 1966 to 1977, Publication No. 38, Reports of the International Symposium on the Computation and Predicted Runoff from Glacers and Glacial Areas, Section of Glaciology of the Soviet Geophysical Committee and Institute of the Geography, Moscow, 138-143, 214-216.

Meier, M.F., Post, A.S., 1962, Recent variations in mass net budgets of glaciers in western North America. International Association of Scientific Hydrology 58, 63-77.

Meierding, T.C., 1982, Late Pleistocene Glacial Equilibrium-Line Altitudes in the Colorado Front Range: A Comparison of Methods, Quaternary Research 18, 289-310.

Menounos, B., Osborn, G., Clague, J.J., Luckman, B.H., 2009, Latest Pleistocene and Holocene glacier fluctuations in western Canada. Quaternary Science Reviews 28, 2049- 2074.

Moberg, A., Sonechkin, D.M., Holmgren, K., Datsenko, N.M., Karlen, W., 2005, Highly variable Northern Hemisphere temperatures reconstructed from low-and high- resolution proxy data, Nature 433:7026, 613-617.

Nesje, A. and Dahl, S.O., 1991, Late Holocene glacier fluctuations in Bevringsdalen, Jostedalsbreen region, western Norway (ca 3200-1400 BP), The Holocene 1, 1-7.

Osborn, G., Menounos, B., Ryane, C., Riedel, J., Clague, J., Koch, J., Clark, D., Scott, K., Davis, P., 2012, Latest Pleistocene glacier fluctuations on Mount Baker, Washington. Quaternary Science Reviews 49, 33-51.

Osipov, E., 2004, Equilibrium-line altitudes on reconstructed LGM glaciers of the northwest Barguzinsky Ridge, Northern Baikal, Russia. Paleogeography, Paleoclimatology, Paleoecology 209, 219-226.

Papineau, J.M., 2001, Wintertime temperature anomalies in Alaska Correlated with the ENSO and PDO. International Journal of Climatology 21, 1577-1592.

Porter S.C., 1976, Pleistocene glaciation in the southern part of the North Cascade Range, Washington. Geological Society of America Bulletin, 87, 61-75.

Porter, S.C., 1977, Present and past glaciation threshold in the Cascade Range, Washington, U.S.A.: Topographic and climatic controls, and paleoclimatic implications. Journal of Glaciology, 18(78), 101-116.

Porter, S.C., 1981, Recent glacier variations and volcanic eruptions. Nature 291, 139- 142.

Porter, S.C., 1986, Pattern and Forcing of Northern Hemisphere Glacier Variations during the Last Millennium. Quaternary Research 26, 27-48.

Putkonen, J. and Michael, O’Neal, 2006, Degradation of unconsolidated Quaternary landforms in the western North America. Geomorphology 75, 408-419.

Reasoner, M.A., Osborn, G., and Rutter, N.W., 1994, Age of the Crowfoot advance in the Canadian Rocky Mountains: A glacial event coeval with Younger Dryas oscillation: Geology 22, 439-442.

Roe, G. and O’neal, M., 2009, The response of glaciers to intrinsic climate variability: observations and models of late-Holocene variations in the Pacific Northwest. Journal of Glaciology 55, 839-854.

Rossignol-Strick, Martine, 1999, The Holocene climatic optimum and pollen records of sapropel in the eastern Mediterranean, 9000-6000BP, Quaternary Science Reviews 18, 515-530

Ruiz, L., Masiokas, M.H., Villalba, R., 2011, Fluctuations of Glaciar Esperanza Norte in the north Patagonian Andes of Argentina during the past 400 yr, Climate of the Past 8, 1079-1090.

Rupper, S., and Roe, G. 2007, Glacier Changes and Regional Climate: A Mass and Energy Balance Approach. Journal of Climate 21, 5384-5401.

Scuderi, L. A., 1987, Glacier variations in the Sierra Nevada, California, as related to a 1200-year tree ring chronology. Quaternary Research 27, 220-231.

Schimmelpfennig, I., Schaefer, J.M., Akcar, N., Ivy-Ochs, S., Finkel, R.C., Schluchter, C., 2012, Holocene glacier culminations in the Western Alps and their hemispheric relevance, Geology 40, 891-894.

Schaefer, J., Denton, G., Kaplan, M., Putnam, A., Finkel, R., Barrell, D., Andersen, B., Schwartz, R., Mackintosh, A., Chinn, T., and Schluechter, C., 2009, High-frequency Holocene glacier fluctuations in New Zealand differ from the northern signature, Science 324, 622-625

Sharp, R.P., 1960, Pleistocene Glaciation in the Trinity Alps of Northern California, American Journal of Science 258, 305-340.

Sigafoos, R. and Hendricks, E.L., 1972, Recent Activity of Glaciers of Mount Rainier, Washington: Botanical evidence of glacier activity. Geological survey professional paper 387-B.

Skinner, C.N., 2003, Fire history of upper montane and subalpine glacial basins in the Klamath Mountains of northern California. in K.E.M. Galley, R.C. Klinger, and N.G. Sugihara (eds.). Proceedings of Fire Conference 2000: The First National Congress on Fire Ecology, Prevention, and Management. Misc. Publication No. 13, 145-151.

Skinner, C.N.; Taylor, A.H.; Agee, J.K., 2006, Klamath Mountains bioregion. In: Fire in California’s Ecosystems. Edited by N.G. Sugihara, J.W. van Wagtendonk, J. Fites- Kaufman, K.E Shaffer, A.E. Thode. University of California Press, Berkeley. 170- 194.

Stokes, M.A., and Smiley, T.L., 1968, An Introduction to Tree-Ring Dating, University of Arizona Press, Tuscon, AZ.

Smith, D. and Lewis, D., 2007, Dendroglaciology. Glacial Landforms, Tree Rings/Dendrochronology 988-994. 64

Smith, D.J., McCarthy, D.P., Colenutt, M.E., 1995, Little Ice Age glacial activity in Peter Lougheed and Elk Lakes provincial parks, Canadian Rocky Mountains, Can. J. Earth Sci. 32, 579-589.

Solomina, O., Haeberli, W., Kull, C., Wiles, G., 2008, Historical and Holocene glacier- climate variations: General concepts and overview. Global and Planetary Change 60, 1- 9.

Speer, J.H., 2010. Fundamentals of Tree-Ring Research: Arizona, 333p.

Tangborn, W., 1980, Two Models for Estimating Climate-Glacier Relationships in the North Cascades, Washington, U.S.A. Journal of Glaciology. 25, 3-21.

Thompson, D.W.J., and Wallace, J.M., 1998, The Arctic oscillation signature in the wintertime geopotential height and temperature fields, Geophysical Research Letters 25, 1297-1300.

Tornes, I., Rye, N., Nesje, A., 1993, Modern and Little Ice Age equilibrium-line altitudes on outlet valley glaciers from Jostedalsbreen, Western Norway: an evaluation of different approaches to their calculation. Arctic and Alpine Research 25, 106-116.

USGS National Department of Water Resources, California State Climatologist: Climate data, North Coast Hydrologic Region, http://www.water.ca.gov/floodmgmt/hafoo/csc/climate_data/northcoast.cfm.

Wanner, H., Beer, J., Butikofer, J., Crowley, T.J., Cubasch, U., Fluckiger, J., Goosse, H., Grosjean, M., Joos, F., Kaplan, J.O., Kuttel, M., Muller, S.A., Prentice, I.C., Solomina, O., Stocker, T.F., Tarasoc, P., Wagner, M., Widmann, 2008, Mid- to Late Holocene climate change: an overview, Quarternary Science Reviews, 1-38.

Winchester V. and Harrison, S., 2000, Dendrochronology and lichenometry: colonization, growth rates and dating of geomorphological events on the east side of the North Patagonian Icefield, Chile, Geomorphology 34, 181-194.

Winchester, V., 2001, Recent Retreat Glaciar Nef, Chilean Patagonia, Dated by Lichenometry and Dendrochronology. Arctic, Antarctic, and Alpine Research 33, 266- 273.

Zhu, H., Xu, P., Shao, X., Luo, H., 2012, Little Ice Age glacier fluctuations reconstructed for the southeastern Tibetan Plateau using tree rings. Quaternary International 1-5 doi: 10.1016/j.quaint.2012.04.011.

Appendix A. Figures

Figure 1. Location map of the Grizzly Valley in the Klamath Mountains.

Figure 1(a). DEM from Geomap App of the Klamath Mountain region with the locations of the Grizzly Lake Glacier and Mt. Shasta indicated by red lines. (b) 3-D image looking to the south constructed with Geomap App, showing the location of the cirque and moraines (M-1 and M-2) beneath Thompson Peak. The elevated position of the cirque at the head of the valley is unique for this region and is likely a contributing factor for increased local precipitation. (c) Index map of California with the Grizzly Valley cirque indicated with a red dot.

Figure 2. DEM of Grizzly Valley cirque

Figure 2. Digital elevation map of the Grizzly Valley cirque, indicating the shape and location of the modern glacier relative to Thompson peak. Grizzly Lake is positioned to the north of Thompson Peak and roughly 290 meters in elevation below the modern Grizzly Lake Glacier. The approximate locations of M-1, M-2, and Mw moraine complexes are also indicated.

Figure 3. Glacial maxima chronology

Figure 3. Diagram summarizing maximum glacial advances documented in the Sierra Nevada, Cascade Range, and Coast Mountains over the last 800 years. A graph of temperature variations for the Northern Hemisphere is also plotted, which was constructed using data from Crowl (2000), which is based on coral, tree-ring, and ice core proxies. The blue bars represent documented intervals of maximum glacial advance and the gray columns indicate periods of synchronous glacial advance between the regions.

Figure 4. M-1 and M-2 moraine crests

Figure 4. Images of M-1 and M-2 depicting how 50 meter transects were made along moraine crests to carry out clast counts. At 1 meter intervals, the underlying clast or soil was described with clast diameter or approximate soil depth. The larger boulders appearing in the foreground on the moraine pictures above are around 1-1.5 meters in diameter. Notice the clear difference in soil development and tree growth between M-1 and M-2.

Figure 5. Sample site location map

Figure 5. Image of the Grizzly Valley cirque looking towards the south. All mapped moraines are identified in the picture with their respective colors. The dendrochronology and cosmogenic samples are identified. The cosmogenic ages are also provided next to the sample locations. Note that the terminus of M-1 extends into the modern day Grizzly Lake.

Figure 6. Tree species for dendrochronology

Figure 6. Images showing the two tree species from which tree cores were sampled in this investigation. Both tree species are conifers, around 10-15 meters tall and 0.6 meters in diameter, and display distinct annual rings in cores.

Figure 7. Pith correction

Figure 7. Schematic diagram showing how to estimate the number of rings missed by a tree core if the pith of the tree was not cored. This number is referred to as the pith correction (PC) in this investigation, and is primarily determined by looking at the curvature of the innermost rings in the core and taking into account their average width. This schematic diagram shows a core with a pith correction of 7 rings.

Figure 8. Moraine age from dendrochronology

Figure 8. This table explains all factors used in determining moraine age from tree ring counts. The ecesis interval and pith correction have to be added to the actual ring count to establish an accurate age for the landform the tree is growing on, in this case a glacial moraine. These ages establish a minimum age for the moraines, because it is possible the tree started growing while the glacier was still in contact and depositing sediment to the moraine, but the moraine is at least as old as the age determined by this equation.

Figure 9. Skeleton plot diagram

Figure 9. A schematic diagram of how the skeleton plots were created based on tree-ring widths. Years with relatively high precipitation increase tree growth rate and result in a relatively wide ring, while dry years typically have limited growth rates and produce narrow rings. On the skeleton plots, wide rings receive short or no lines, while narrow rings are represented with long lines. Holes were poked in the cores at 10-year intervals, like the one seen in the fifth ring from the right in the image, to help keep track of annual rings.

Figure 10. ELA climate conditions and envelope

Figure 10. Graph from Leonard (1989) with the winter accumulation and summer temperature at the equilibrium lines of 32 glaciers worldwide plotted. The plotted points generally fall within the climate envelope, which is defined in equation 1 and developed by Kotlyakov and Krenke (1982), which gives parameters for typical climate conditions at the ELA of modern glaciers.

Figure 11. Location map of temperature stations

Figure 11. Shaded relief model indicating the locations of the weather stations used to construct the local temperature lapse rate. Our field area is indicated with the yellow star. The weather stations were located at various altitudes, which formed a clear linear relationship of mean summer temperature.

Figure 12. Location map of precipitation stations

Figure 12. Digital elevation map displaying the locations of weather stations used for constructing a local precipitation. These weather stations were chosen, all east of the study area, due to their proximity to the study area and to minimize the effect of the regional rain shadow effect.

Figure 13. M-1 and M-2 crest comparison

Figure 13. Modified images showing the difference in moraine crest degradation between the M-1 and M-2 moraines. As indicated in the figure, the M-1 moraine has a broad crest and the M-2 moraine has a sharp crest, indicative of recent deposition. The difference in soil development on the moraines is also clearly seen in the pictures.

Figure 14. M-2 profile

Figure 14. Schematic diagram of the M-2 profile, with descriptions of sections of variable compositions. The picture of the moraine side proximal to the glacier shows substantial rock flour at the moraine surface, while the surface of the distal side has a greater ratio of boulders to rock flour. The rock flour on the proximal side of the moraine is intuitive because it is produced by sub-glacial erosion, after which it is incorporated into the near side of the moraine.

Figure 15. M-1 soil development

Figure 15. Modified image indicating the three distinct soil horizons, on the surface of M- 1. The descriptions of these soil horizons are summarized in Table 3. These soil horizons clearly distinguish M-1 from M-2, which has not had sufficient time to develop any of these soils. Notice the hammer handle for scale.

Figure 16. Moraine surface soil/clast diagrams

Figure 16. Scatter plots of data collected during the 50 meter clast count transects. The plots indicate a significantly higher soil to clast ratio on the surface of M-1 in comparison to M-2. Only 3 counts (from 1 meter intervals) were soil on M-2, highlighting the recent moraine deposition relative to M-1.

Figure 17. 10Be cosmogenic ages

Figure 17. Graph showing the ages calculated from 10Be results in the Grizzly Valley cirque. Six of the samples reflect significant inheritance, when compared with the relative and tree-ring ages of the moraines (M-1 and M-2). However, the ages with inheritance do overlap with documented glacial advances in the Sierra Nevada, as indicated by the blue rectangles.

Figure 18. Grizzly Valley cirque cosmogenic sample map and ages

Figure 18. Digitized image indicating 10Be cosmogenic sample sites and their respective ages. The picture is taken from the eastern side of the cirque, looking towards the southwest. Thompson Peak is just out of view in the top-left corner of the image. All samples indicated were taken from boulders, with the exception of the blue dot (5206 ka) sample site, which was collected from striated bedrock. Also indicated on the figure is the outlines of M-1, M-2, and Mw.

Figure 19. Tree ages with ecesis interval

!!M1:!652(716!

!M2:!103(167!

Figure 19. Diagram showing the plotted tree ages growing on the M-1 and M-2 moraines. The oldest tree age (used for moraine age determination) is highlighted in yellow for each moraine and the blue bar indicates the moraine age after the addition of the ecesis interval range. The M-2 moraine was deposited at least 103-167 ybp and M-1 was deposited a minimum of 652-716 ybp.

Figure 20. Graphs of the M-1 standardized tree-ring chronologies

a.)

b.)

c.)

out of sync

Figure 20. Graphs of the standardized (index value) chronologies for the east and west limbs of the M-1 moraine a.) and b.), respectively. The curves reflect deviations from mean ring width through time, which is related to environmental conditions (primarily precipitation) during tree growth. The chart indicated by c.) shows the comparison between the two standardized chronologies, with periods that appear out sync indicated by the yellow polygon.

Figure 21. Tree-ring chronology correlations

Figure 21. Figure depicting the correlation between tree-ring width data from Mt. Eddy (Graumlich) and Fryday Ridge with that of the east and west limbs of M-1, as well as river flow data from the Trinity River in Weaverville (USGS - National Water Information System). The tree-ring width data is graphed as a 7 year running mean of annual tree-ring widths. The Trinity River flow represents a 7 year running mean of flow data (ft2/sec) from October-April. The correlation makes it possible to assume historical tree-ring widths accurately represent climate conditions during tree growth. The graph also shows the overlap of the regional wet episode in southern Oregon, as documented by Keen (1937), with a wet period identified in the tree-rings in Grizzly Valley. The depositional age range of M-2 is also shown in relation to the climate record, which were used together to better constrain the timing M-2 deposition. Also notice the significantly wide tree ring associated with the 1983 El Nino.

Figure 22. DEM of cirque with ELAs

Figure 22. Three digital elevation models with the reconstructed glacier bodies during the deposition of M-1, M-2, and the modern glacier. The ELA for each glacier extent is represented by the black line on the glacier, separating the accumulation zone from the ablation zone. The glacier area of the accumulation zone (above the ELA) make up 65% of the total glacier area. The altitudes of the reconstructed and modern equilibrium-lines are summarized in the table in the top right of the figure. The change in ELA for M-1 and M-2 is relative to the ELA of the modern glacier. These ELAs were used in calculating the temperature and precipitation changes during M-1 and M-2 glacier extent. The ELAs for each glacial extent are indicated with the thick bold line running laterally across the glacier.

Figure 23. Climate estimates at ELAs

Figure 23. Diagrams showing the mean summer temperature depression and winter accumulation increase needed to shift the ELAs of the modern glacier to the paleo-ELA conditions. Plotted on a plot of mean summer temperature and winter accumulation conditions at the ELA of 32 glaciers worldwide, created by Leonard (1989).

100

Figure 24. Location map of proximal tree-ring studies

Figure 24. Location map of the nearby sites where dendrochronology investigations have been carried out by other researchers, Mt. Eddy (Graumlich, 1981) and Fryday Ridge (Briffa and Schweingruber, 1983). Both of these tree-ring studies showed tree-ring index values that correlated to those in our investigation, in the Grizzly Valley cirque. The data for these tree-ring studies is publicly available on the NOAA Paleoclimate website. 101

Figure 25. Temperature records

102

Figure 25. Figure displaying Northern Hemisphere temperature data over the last 1000 years. The Crowley and Lowery (2000) temperature record is derived from tree-ring, ice core, and historical data presented as a decadally smoothed time series. The Moberg et al. (2005) reconstruction is based on sedimentation and tree-rings, presented as deviation from the 1961-1990 mean. The Huang (2004) temperature data extends back ~500 ybp and is derived from multiple proxies, including borehole temperatures, 20th century meteorological records, and multi-proxy paleoclimate records. The Huang (2004) temperature data is depicted in the figure as a decadally smoothed time series. The global temperature reconstruction (Leclercq and Oerlemans, 2012) is based on the glacier length records of 308 glaciers by using a linear response equation to calculate temperature from glacier length. The data is presented as the degrees C anomaly with respect to the 1961- 1990 average. Ranges of documented glacial maxima in the Sierra Nevada and Cascade Ranges are also shown. Periods of increased precipitation in the Grizzly Valley, based on dendrochronology in this study, are indicated with the tan boxes.

103

Figure 26. Regional subdivisions for dendrochronology

Figure 26. Figure modified from Graumlich (1987), showing the separate regions corresponding to the various tree-ring investigations corresponding to figures 28 and 29. The study area is indicated by the red circle, located in the region referred to as the Southern Valley by Graumlich (1987).

104

Figure 27. Precipitation records in NW North America

105

Figure 27. Figure modified from Graumlich (1987) and corresponding to Fig. 26. The boxes outlined in a red dashed line indicate time intervals when precipitation patterns in the Southern Valley (the field study region) are out of synch with the Western Basin and Columbia Lowlands. The box labeled M-2 indicates a relatively wet period in our field area (1850-1890), which corresponds to the precipitation patterns reflected in the tree- ring widths from our dendrochronology investigation. During this period, the Grizzly Valley glacier advanced to its M-2 extent, while droughts characterized the climate conditions immediately to the north (Columbia Lowlands and Western Valley).

106

Figure 28. Precipitation records of study region

Figure 28. A graph modified from Graumlich (1987) and corresponding to Fig. 26. The beige box highlights the wet episode from 1850-1895. The red box indicates a relatively wet period from 1850-1870, during which a significant increase in precipitation occurred. Our tree-ring records also indicate this period was wet and suggest the glacier entered a state of positive mass balance during this period. The Sierra Nevada and Cascade ranges appear to have reached their LIA glacial maxima around 1850, and retreated shortly thereafter. This graph and our tree-ring records suggest the Grizzly Valley Glacier sustained its mass until around 1870 C.E.

107

Figure 29. Precipitation lapse rate

Figure 29. Graph showing the mean snow water equivalent (cm) of winter accumulation from 1970-1996 recorded on April 1. The measurements were made at four weather stations proximal to the field area at different elevations. A power regression curve is fit through the data points and extrapolated to project precipitation at higher elevations. The weather stations used for the precipitation lapse rate are Bear Basin, Big Flat, Red Rock Mountain, and Whalen.

108

Figure 30. Temperature lapse rates

Figure 30. Local temperature linear lapse rate calculated from mean summer temperature June – September (1971-2000) from nine local winter stations at various altitudes. Local lapse rates calculated for the central Sierra Nevada (1Bowerman and Clark, 2011) and for the Three Sisters Volcanoes region (2Marcott, 2005) are also plotted for comparison. A linear regression is fit to the data to establish mean summer temperatures at the modern and paleo-ELAs.

109

Figure 31. Tree-ring widths and ENSO record

Tree-Ring Width Correlation with ENSO (D’Arrigo et al., 2005)

Dev. from mean

ENSO (D’Arrigo et al., 2005) M-1 tree-ring widths

Figure 31. The El Nino-Southern Oscillation record is based on tree-ring chronologies and instrumental records in southwestern Mexico (D’Arrigo et al., 2005). The M-1 tree- ring chronology curve represents the averaged index values for the east and west limbs of M-1.

110

Figure 32. Reconstructions of the Grizzly Valley glacier

111

Figure 32. Panoramic photo taken from the north side of Grizzly Lake, looking south towards Thompson Peak. The approximate glacier limits during the deposition of M-1 and M-2 are digitized, as well as the modern day glacier extent.

112

Appendix B. Tables

Table 1. Documented ecesis intervals in North America

Table 1. Table summarizing documented ecesis intervals for trees growing on moraines in western North America. These ages range from 0-93, which is a function of local environmental conditions and proximity to a seed source.

113

Table 2. Summary of boulder and moraine characteristics

Table 2. Table summarizing the boulder and moraine characteristics documented in the field. These results were taken into account when applying relative ages to the M-1 and M-2 moraines, as well as comparing these characteristics with other documented LIA moraines.

114

Table 3. Summary of M-1 soil descriptions

Table 3. Summary of M-1 surface descriptions. These descriptions correspond to the soil horizons seen in Figure 15.

115

Table 4. Temperature and winter precipitation changes at calculated ELA’s

Table 4. Summary of the estimated summer temperature and winter accumulation conditions at the ELAs of the modern glacier and the paleo-glaciers M-1 and M-2. These values were calculated from local lapse rates.

116

Table 5. COFECHA statistical analysis values

Site # cores # rings mean core length (mm) st. dev. length mean ring width (mm) st. dev width intercorrelation M-1 East 10 3284 328.4 110.9 0.725 0.273 0.276 M-1 West 8 2693 319.3 87.4 0.716 0.262 0.285 Mt. Eddy 23 10655 463.3 75.4 0.775 0.215 0.484 Fryday Ridge 24 10444 462 54.7 0.777 0.264 0.402

Table 5. Chart displaying values derived from the COFECHA statistical analysis of the M-1 east and west limb series as well as those for the Mt. Eddy and Fryday Ridge. For each chronology, the table shows the total number of cores, number of rings, mean core length (mm), standard deviation of the length, mean ring width (mm), standard deviation of the ring widths, and the series intercorrelation value.

117

Table 6. Descriptions of precipitation data weather stations

Table 6. Showing the coordinates, elevation (m), and winter accumulation (cm SWE) of the four weather stations used to construct the local precipitation lapse rate. These locations are indicated on Figure 12.

118

Appendix C. Lawrence Livermore Data

21.3 20.5 1.685E-14 5.252E-14 1.295E-14 1.375E-14 1.155E-14 1.239E-14 1.226E-14 6.639E-14 4.507E-14 6.108E-15 7.778E-15 4.486E-15 6.446E-15 1.072E-14 1.314E-14 1.640E-14 5.070E-15 6.723E-16 ERROR BKGDS) 24.5 21.6 21.3 23.4 FOR 2.892E-12 2.912E-12 2.837E-12 2.859E-12 2.827E-12 2.832E-12 2.844E-12 8.504E-12 8.557E-12 9.780E-13 9.627E-13 5.302E-13 5.237E-13 4.742E-13 5.501E-13 8.362E-13 1.926E-13 4.868E-15 (CORR. (CORR. 23.3 21.0 20.9 23.4 24.5 21.2 10Be/9Be RATIO RATIO 22.8 22.5 20.5 22.7 24.0 22.5 6.72E-16 6.72E-16 6.72E-16 6.72E-16 bkgd_error BKGD)

23.1 22.6 21.1 22.1 23.2 22.4 4.87E-15 4.87E-15 4.87E-15 4.87E-15 bkgd_ratio

(SAMPLE 10Be/9Be RATIO 23.2 22.7 20.8 23.2 22.8 23.6

BORON) 1.685E-14 5.252E-14 1.295E-14 1.375E-14 1.155E-14 1.239E-14 1.226E-14 6.639E-14 4.507E-14 6.108E-15 7.778E-15 4.486E-15 6.446E-15 1.070E-14 1.312E-14 1.639E-14 5.025E-15 6.723E-16 ratio_err1 RATIO FOR

21.4 23.5 22.7 20.5 23.5 24.4 23.1 22.8

Be_ratio1 10Be/9Be 10Be/9Be 2.892E-12 2.912E-12 2.837E-12 2.859E-12 2.827E-12 2.832E-12 2.844E-12 8.504E-12 8.557E-12 9.780E-13 9.627E-13 5.302E-13 5.237E-13 4.791E-13 5.550E-13 8.411E-13 1.975E-13 4.868E-15

CORRECTED (

21.9 24.2 24.6 23.0 24.1 24.2 22.4 23.1 0.999944

Truefrac 0.9999329 0.9999432 0.9999324 0.9999335 0.9999319 0.9999324 0.9999334 0.9999432 0.9998542 0.9998444 0.9998704 0.9998393 0.9948348 0.9968809 0.9972123 0.9846585 0.6847532

22.3 24.0 24.7 23.4 25.5 24.4 25.7 23.8

0.002082 0.004603 0.0059117 0.0184271 0.0045438 0.0048251 0.0040522 0.0039854 0.0038133 0.0232948 0.0158133 0.0016943 0.0015742 0.0022617 0.0037532 0.0025393 0.0016928 0.0002278 exterror 23.6 25.6 24.3 23.6 25.8 24.2 24.8 24.2

0.004349 0.002143 0.0041808 0.0113343 0.0038666 0.0038719 0.0040354 0.0043033 0.0187369 0.0152772 0.0027292 0.0013855 0.0017734 0.0032623 0.0037687 0.0057496 0.0017633 0.0002359 interror 23.3 21.8 23.0 26.8 23.7 23.1 25.4 26.8 25.1 24.6 22.9

1.014652 1.021655 1.002983 2.983979 3.002298 0.9952785 0.9920543 0.9938175 0.9980444 0.3431422 0.3377822 0.1860217 0.1837516 0.1681216 0.1947507 0.2951156 0.0693134 0.0017079 r_to_rstd 25.2 22.6 15.9 22.6 25.1 23.7 22.1 25.6 26.0 24.3 27.1 23.7 2 4 6 6 6 3 3 3 3 2 15 16 16 15 14 14 10 10

runs 23.9 23.3 24.6 23.3 24.2 25.2 24.2 25.1 25.6 26.5 26.7 25.7 07KNSTD3110

# CAMS 2.85E-12 7.6 6.4 3.7

(5±1)x10^-5 BE34240 BE34241 BE34242 BE34243 BE34244 BE34245 BE34246 BE34226 BE34227 BE34154 BE34219 BE34212 BE34213 BE34260 BE34262 BE34263 BE34264 BE34291 14.1 24.5 23.3 25.7 26.6 26.6 25.9 25.7 24.0 24.9 23.6 23.6 26.6 24.6 26.7 27.2 25.7

7.9 8.1 4.4 11.2 14.7 24.3 23.6 26.8 27.0 26.4 26.8 25.4 25.8 25.2 24.6 24.0 23.8 26.7 26.6 27.1 27.6 26.8 NAME (microA) SAMPLE SAMPLE KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD 9422 KNSTD 9422 KNSTD 1032 KNSTD 1032 KNSTD549 KNSTD549 Gc 08-08-3 (cirque lip) 08-06-2GR (brk in M2) GC 08-08-1 (M1) GC 08-11-1 (M1) DH blank 9.4 8.4 6.1 6.4 = = 12.7 24.2 24.0 26.8 27.6 26.8 26.0 26.8 27.9 25.9 24.2 23.5 25.1 16.3 24.7 26.7 26.0 27.9 stds= for normalization:

Current factor for standard for ID DATE used background background correction ratio Laboratory BE34291 BE34263 BE34260 BE34264 BE34262 BE34154 BE34219 BE34265 BE34266 BE34267 BE34268 BE34226 BE34227 BE34240 BE34241 BE34242 BE34243 BE34244 BE34245 BE34246 BE34212 BE34213 8 November 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8 2012 November 8

Standard 10/9 Carrier Boron National Sample

Livermore

Lawrence STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD STANDARD Heermance Heermance Heermance Heermance Heermance DH blank GC 08-08-1 (M1) Gc 08-08-3 (cirque lip) GC 08-11-1 (M1) 08-06-2GR (brk in M2) KNSTD 1032 KNSTD 1032 KNSTD 549 KNSTD 549 KNSTD 549 KNSTD 549 KNSTD 9422 KNSTD 9422 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD3110 KNSTD549 KNSTD549 a.) b.)

119 a.) Lawrence Livermore National Laboratory Grizzly Valley cirque results for two M-1 boulder samples and two striated bedrock samples. Results also show the AMS results for the standards and blanks run with our samples. b.) Current measurements during AMS data collection.

120

13.4

13.8 18.7

BKGDS)

FOR FOR

2.878E-12

1.707E-14

1.772E-11

7.776E-14

6.147E-15

1.611E-14

1.946E-13

5.066E-13

8.492E-12

8.475E-12

9.712E-13

9.783E-13

5.285E-13

5.290E-13

2.851E-12

2.836E-12

2.852E-12

2.849E-12

2.853E-12

2.833E-12 2.870E-12

13.8

19.2

14.5 16.9

(CORR. (CORR.

RATIO

10Be/9Be RATIO

14.3

19.2 16.4

19.5

1.39E-15

1.39E-15 1.39E-15

bkgd_error

15.3

16.5

18.6 20.1

BKGD)

17.0

20.9

17.0

19.2

19.4

19.7

1.71E-14

bkgd_ratio

(SAMPLE (SAMPLE

10Be/9Be RATIO

BORON)

4.653E-14

1.377E-15

1.544E-13

3.639E-15

1.258E-15

1.661E-15

3.977E-15

9.792E-15

3.383E-14

3.691E-14

6.649E-15

8.155E-15

4.094E-15

3.906E-15

1.980E-14

1.101E-14

1.289E-14

1.324E-14

1.143E-14

1.429E-14

1.174E-14

ratio_err1

16.9

21.6

17.9

19.7

20.1

18.9 RATIO FOR FOR

19.7

18.1

22.0

20.1

19.6

20.6

19.4

10Be/9Be 10Be/9Be

2.878E-12

1.707E-14

1.774E-11

9.483E-14

2.321E-14

3.318E-14

2.117E-13

5.237E-13

8.492E-12

8.475E-12

9.712E-13

9.783E-13

5.285E-13

5.290E-13

2.851E-12

2.836E-12

2.852E-12

2.849E-12

2.853E-12 2.833E-12

2.870E-12

CORRECTED

BE_ratio1 ( (

19.2

19.1

16.6

21.9

19.8

19.1 20.9

19.4

0.99994

0.975302 0.999914

Truefrac

0.9999403

0.8089737

0.9998331

0.9171783

0.9408037

0.9950993

0.9948553

0.9999459

0.9999449

0.9999116

0.9999101

0.9998956

0.9998987

0.9999325

0.9999407

0.9999401

0.9999381 0.9999333

18.7

19.6

21.1

19.1

23.0

20.3

20.4

19.7

20.5

19.2

17.8

15.4 16.4

0.004029

0.004644

0.0001742

0.0342731

0.0004493

0.0000299

0.0005827

0.0013953

0.0022215

0.0109334

0.0121238

0.0018832

0.0028614

0.0012299

0.0009371

0.0069478

0.0033793

0.0037096 0.0050153

0.0041187

exterror

20.8

21.7

20.7

19.7

22.2

19.9

20.4

20.6

21.3

19.1

17.9 16.2 15.8

0.054164

0.000484

0.003864

21.8

22.4

20.4

19.4

21.0

16.6

20.6

21.8

23.4

21.2

19.7

18.2

16.2

0.0163251

0.0004833

0.0012768

0.0004415

0.0012011

0.0034357

0.0118689

0.0129521

0.0023329

0.0023702

0.0014365

0.0013707

0.0054272

0.0045239

0.0044478

0.0040117 0.0038645 0.0039168 interror

19.6

23.7

20.7

19.7

19.2

20.0

24.0

22.7

23.3

18.6

20.9

16.8

18.2

1.00028

1.009789

6.223697

0.008145

2.979555

2.973826

1.000717

1.001058

0.993968

1.006843

0.0059889

0.0332731

0.0116406

0.0742848

0.1837427

0.3407624

0.3432785

0.1854232

0.1856198 0.9951901 0.9996482

r_to_rstd

runs

15.1

23.8

21.3

20.1

25.9

22.2

23.6

24.8

24.3

20.1

21.2

16.1 19.7

16.2 #

07KNSTD3110

CAMS CAMS

2.85E-12

BE34765

BE34576

BE34575

BE34572

BE34571

BE34570

BE34569

BE34568

BE34791

BE34790

BE34751

BE34750

BE34743

BE34742

BE34764

BE34763

BE34762

BE34761

BE34760

BE34759

BE34758

(5±0.9)x10^-5

4.5

7.4

8.8

5.5

6.8

5.1

21.6

25.2

20.9

20.1

19.5

22.5

20.3

24.4

22.8

21.0

22.2

18.3

19.6

16.6 4.0

8.2

4.9

4.3

5.7

22.2

23.4

19.4

16.9

21.4

22.9

21.5

24.4

25.7

21.4

20.0

18.8

19.9

10.9 16.0

NAME

3.8

9.1

4.7 2.5

4.2

26.5

25.1

19.9

19.1

15.4

27.2

21.0

24.8

22.5

13.8

24.8

23.4

21.6

20.7

10.5

12.6

stds=

KNSTD3110

1302blank

CRONUS-A

outer outer M-2 (3)

outer outer M-2 (2)

outer outer M-2 (1)

inner M-2 (2)

inner M-2 (1)

KNSTD 9422 KNSTD

KNSTD 1032 KNSTD

KNSTD549

KNSTD3110

KNSTD3110 SAMPLE SAMPLE

(microA) resutlts

for for

normalization:

factor factor

for for standard standard

Current

for for used

Laboratory

DATE

background background

correction ratio ratio

8 February 2013 February 8

Boron Boron

Carrier

10/9 10/9 Standard Standard

National

BE34743

BE34742

BE34765

BE34764

BE34763

BE34762

BE34761

BE34760

BE34759

BE34758

BE34791

BE34790

BE34751

BE34750

outer M-2 (3) M-2 outer

outer M-2 (2) M-2 outer

outer M-2 (1) M-2 outer

inner M-2 (2) M-2 inner

inner M-2 (1) M-2 inner BE34575 BE34576

Livermore Sample Sample

Standard

Heermance

STANDARD

Lawrence

KNSTD549

KNSTD3110

KNSTD 9422 KNSTD

KNSTD 1032 KNSTD

GC 08-12-01 GC

GC 08-11-03 GC

GC 08-11-02 GC

GC 08-10-03 GC

GC 08-10-02 GC

CRONUS-A 1302blank c.) d.)

121 c.) Lawrence Livermore National Laboratory Grizzly Valley cirque results for two inner and three outer M-2 boulder samples. Results also show the AMS results for the standards and blanks run with our samples. d.) Current measurements during AMS data collection.

122

Appendix D. COFECHA Output Data

M-1 East Limb Index Values

!"#$ %& !"#$ %& !"#$ %& !"#$ %& !"#$ %& !"#$ %& !"#$ %& '()) )*+, '(-- .)*)( '(/, )*0' '/+- .)*+1 '/(, '*-+ ',1- .'*12 ',2, )*)' '()' .)*-0 '(-0 .)*2+ '(,) .)*)+ '/+0 .)*0- '//) .)*'/ ',10 .)*0- ',() )*21 '()1 .)*() '(-2 )*+0 '(,' .)*'1 '/+2 .)*12 '//' .)*'( ',12 )*/- ',(' .)*0) '()+ .)*'+ '(-( .)*1, '(,1 .)*)( '/+( )*2( '//1 )*1( ',1( )*++ ',(1 .)*/- '()- )*), '(-/ .)*)0 '(,+ )*-' '/+/ )*'' '//+ )*(( ',1/ .)*)1 ',(+ )*-+ '()0 .)*)+ '(-, )*/1 '(,- )*/- '/+, )*/, '//- .)*-- ',1, )*/, ',(- .)*'2 '()2 '*'+ '(0) .)*-) '(,0 )*') '/-) .)*/2 '//0 .)*2' ',+) .)*+/ ',(0 )*(' '()( )*10 '(0' )*-, '(,2 .)*+, '/-' .)*1- '//2 .)*)2 ',+' .)*+2 ',(2 '*'( '()/ .)*1+ '(01 .)*20 '(,( .)*)2 '/-1 .'*)0 '//( )*+) ',+1 .)*(+ ',(( .)*)' '(), .)*0) '(0+ .'*,+ '(,/ .)*') '/-+ )*-, '/// )*)/ ',++ .'*11 ',(/ .)*/0 '(') .)*'( '(0- .)*), '(,, .)*'1 '/-- )*+/ '//, )*/( ',+- .)*)- ',(, )*00 '('' )*(2 '(00 .)*)/ '/)) .)*,0 '/-0 .)*-/ '/,) .'*/( ',+0 )*/2 ',/) '*)) '('1 .)*)/ '(02 )*0+ '/)' .)*,+ '/-2 '*)2 '/,' )*)2 ',+2 )*2- ',/' )*-' '('+ )*0+ '(0( .)*)+ '/)1 .)*,1 '/-( .)*1, '/,1 )*2) ',+( .'*)+ ',/1 .)*0' '('- .)*1/ '(0/ )*-1 '/)+ )*-+ '/-/ )*,1 '/,+ .)*(1 ',+/ .)*12 ',/+ .'*-+ '('0 .)*+' '(0, )*/, '/)- )*-' '/-, )*)+ '/,- .)*11 ',+, )*-+ ',/- )*,2 '('2 .'*01 '(2) .)*), '/)0 )*2+ '/0) .)*2, '/,0 .)*1( ',-) )*01 ',/0 )*0/ '('( )*1+ '(2' .)*', '/)2 )*(' '/0' )*++ '/,2 )*)+ ',-' .)*(0 ',/2 .)*-, '('/ )*+( '(21 .)*11 '/)( )*(1 '/01 )*(2 '/,( )*/) ',-1 )*10 ',/( )*/0 '(', )*20 '(2+ )*0' '/)/ )*+) '/0+ .)*/2 '/,/ )*/0 ',-+ )*,) ',// .)*1( '(1) )*(, '(2- .'*2+ '/), .)*01 '/0- )*++ '/,, .)*/0 ',-- )*1) ',/, .)*'+ '(1' .)*1, '(20 )*(( '/') .)*/0 '/00 )*-1 ',)) .'*+- ',-0 )*') ',,) )*/, '(11 )*/, '(22 .)*++ '/'' .'*11 '/02 .'*)1 ',)' .)*'( ',-2 .)*++ ',,' .)*12 '(1+ .)*)/ '(2( .)*-' '/'1 )*-/ '/0( .'*') ',)1 )*)- ',-( )*)) ',,1 .)*02 '(1- .'*'1 '(2/ )*-, '/'+ )*0( '/0/ )*-0 ',)+ .)*', ',-/ )*(/ ',,+ .'*0) '(10 )*22 '(2, .)*12 '/'- )*1' '/0, )*++ ',)- )*/+ ',-, )*1+ ',,- )*0' '(12 )*-( '(() )*/+ '/'0 .)*)- '/2) )*+0 ',)0 '*11 ',0) )*21 ',,0 .1*,, '(1( )*', '((' )*+1 '/'2 )*-2 '/2' .)*)- ',)2 )*)+ ',0' .)*+, ',,2 '*+2 '(1/ .)*02 '((1 .)*1( '/'( )*-) '/21 .'*0' ',)( .)*'2 ',01 .)*0- ',,( )*01 '(1, )*') '((+ )*-- '/'/ .)*2( '/2+ )*'1 ',)/ )*(/ ',0+ .)*/+ ',,/ .'*0/ '(+) .)*'( '((- .)*'' '/', .)*() '/2- )*,, ',), .'*0- ',0- )*'/ ',,, )*++ '(+' .'*)1 '((0 )*1' '/1) .)*2) '/20 )*0/ ',') '*)+ ',00 )*,1 1))) )*(, '(+1 .)*11 '((2 )*') '/1' )*1( '/22 .)*-- ','' )*() ',02 .)*1( 1))' )*/- '(++ )*/) '((( .)*') '/11 )*,+ '/2( )*-, ','1 .)*-2 ',0( )*/- 1))1 )*1' '(+- )*-1 '((/ )*1/ '/1+ )*0, '/2/ )*-2 ','+ )*11 ',0/ )*), 1))+ )*)- '(+0 )*01 '((, )*(+ '/1- .)*'1 '/2, '*'( ','- )*0' ',0, )*2- 1))- )*-/ '(+2 )*+0 '(/) )*22 '/10 .1*)/ '/() )*,1 ','0 )*)) ',2) )*2+ 1))0 )*'/ '(+( )*/) '(/' )*2' '/12 .)*+( '/(' .)*2( ','2 )*02 ',2' .'*'/ 1))2 )*-) '(+/ .)*(0 '(/1 .)*-' '/1( )*0) '/(1 .)*(0 ','( )*01 ',21 .'*1/ 1))( '*+1 '(+, .'*1' '(/+ .)*0, '/1/ )*'( '/(+ )*+, ','/ .)*/1 ',2+ .)*1) 1))/ )*'- '(-) )*0/ '(/- .'*)0 '/1, )*)2 '/(- .)*,( ',', .'*)2 ',2- )*-( 1)), .)*1) '(-' .)*'/ '(/0 .'*'2 '/+) .)*+/ '/(0 )*), ',1) .)*-1 ',20 .)*-( 1)') .+*)' '(-1 )*++ '(/2 )*'/ '/+' .)*)' '/(2 .'*-2 ',1' .)*+1 ',22 )*'( 1)'' )*22 '(-+ )*'0 '(/( )*20 '/+1 .)*'0 '/(( )*1, ',11 )*0) ',2( )*11 '(// )*2) '/++ )*+' '/(/ )*++ ',1+ .)*)0 ',2/ .)*+)

123

M-1 West Limb Index Values

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

124