Holocene Glacial and Paleoclimate Reconstructions for the Lyman Glacier and the North Cascades, Washington

Holocene glacial and paleoclimate reconstructions for the Lyman Glacier and the North Cascades, Washington

Thesis Proposal for the Master of Science Degree, Department of Geology, Western Washington University, Bellingham, Washington

Harold N. Wershow July 14, 2014

Approved by Advisory Committee Members:

Dr. Douglas Clark, Thesis Committee Chair

Dr. Robert Mitchell, Thesis Committee Advisor

Dr. Scott Linneman, Thesis Committee Advisor

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Introduction The glaciers of the North Cascades are a vital part of the alpine environment and are a critical water supply for fisheries, agriculture, industry and people throughout the region. The region’s aquatic and terrestrial ecosystems depend upon on the sustained flow of glacial melt water during the otherwise dry Cascadian summers (Riedel and Larrabee, 2011). Human users also depend upon the summer availability to supplement reservoirs and groundwater. The glacial water supply’s local significance is evident from Chelan County’s interest in the status of glaciers that contribute meltwater to Lake Chelan (pers. comm., Scott Buehn, Water Resource Engineer, Chelan County Public Utility District). To properly manage their water resources, users like Chelan County need to better understand how the glacial supply has changed in the past, and will change in the future. For these reasons, prominent glacier monitoring campaigns are being conducted in the North Cascades by the National Park Service, the U.S. Geological Survey, and the North Cascades Glacier Climate Project (Pelto and Riedel, 2001). This information informs year-to-year melt water projections, but it does not give a long-term perspective on changing glacial water supplies. The histories of alpine glaciers provide insight into how and why glacial water supplies change over time. Alpine glaciers are sensitive indicators of climate change; they advance and retreat (fluctuate) rapidly in response to variations in temperature and precipitation (e.g., Leonard, 1989). As such, glacial fluctuation histories reflect past changes in climate, and they help constrain forecasts of future changes in glacial extent. Despite these benefits, the Holocene history of glacial fluctuations in the North Cascades is not well understood. Existing Holocene glacial and paleoclimate records are generally poorly dated, discontinuous, spatially limited, or a combination of the three. The foremost goal of my project is to develop the first high-resolution, continuous history of Holocene fluctuations and climate for a representative glacier in the North Cascades. The North Cascades’ Lyman Glacier (Figure 1) is well suited to refining the regional Holocene glacial and climate record for several reasons: 1) it is a small, discrete alpine glacier with a well-preserved Holocene moraine record; 2) the rock flour transported in its meltwater is deposited in a series of bedrock-dammed lakes below the outermost late- Holocene moraine, preserving a continuous lake sediment record (Dahl et al., 2003); 3) historical documentation of the glacier began as early as the late 1890s with photographs and terminus measurements; 4) an exceptional modern climate record exists for the site, including snowpack data since 1926 and meteorological data since 1979 from a nearby snow telemetry site (SNOTEL); and 5) the Lyman Glacier has an annual mass-balance record dating back to 1986 (Pelto and Riedel, 2001). Crucially, the Lyman Glacier’s fluctuation record is of regional significance because North Cascades glaciers tend to vary in concert (Pelto and Riedel, 2001). Glacial fluctuation proxy records vary considerably in their availability and utility, and as such they are most informative when multiple records are available to complement each other (Osborn et al., 2012). At the Lyman Glacier, proxy records from moraines, lake sediments, historical documentation, climate data, and mass-balance measurements allow development of a continuous and quantified multi-proxy record of Holocene glacial fluctuations. The glacial fluctuation record, in turn, can be linked to paleoclimate conditions via glacial equilibrium-line altitudes (ELAs). These relationships will allow me to develop a continuous Holocene climate record (e.g., Bowerman and Clark, 2011).

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This record will be informative to users of the glacial water supply as well as scientists who work in the North Cascades.

Background

Holocene Climate and Glacial Fluctuations in the western North American Cordillera

The end of the Last Glacial Maximum (LGM) in North America began with wide-spread retreat of mountain glaciers around 19-20 ka (Clark et al., 2009). Post-LGM warming in the western North American Cordillera, and associated glacier retreat, peaked around 13- 15 ka (Marcott, 2011) This warming trend was reversed in the late Pleistocene, when many mountain glaciers re-advanced (Davis et al., 2009). Many, but not all, of these advances have been correlated with the European Younger Dryas cold event from 12.9 – 11.6 ka (Davis et al., 2009). The onset of the Holocene at 11.7 ka saw a resumption of the post-LGM warming trend (Walker et al., 2008). Globally, temperatures increased ~ 0.6 ˚C in the early Holocene, exhibited little change from ~ 9.5 ka. to ~ 5.5 ka., and then gradually dropped ~ 0.7 ˚C through the mid- to late Holocene, culminating in the Little Ice Age (Marcott et al., 2013). Glaciers in the western North American Cordillera (Figure 2) followed suit, retreating during the Altithermal (earliest Holocene) and then advancing during the Neoglacial of the mid- to late-Holocene (Davis et al., 2009). The Little Ice Age (LIA), which began in some places as early as 1000 years ago (Menounos et al., 2009), saw these glaciers reach their maximum Holocene extents (Menounos et al., 2013). Unfortunately, in many locales LIA advances over-rode the deposits left behind by earlier Neoglacial advances, significantly obscuring the Neoglacial record. Although this basic pattern repeats itself throughout the Cordillera, regional histories differ in timing and magnitude (Figure 3). Furthermore, there is considerable uncertainty regarding potential early Holocene glacial advances, the timing and magnitude of multiple Neoglacial advances, and the onset of the Little Ice Age.

WESTERN CANADIAN CORDILLERA The western Canadian Cordillera is the most extensively studied region in the western North American Cordillera (e.g., Davis et al., 2009; Menounos et al., 2013, 2009; Osborn et al., 2007). The latest Pleistocene saw glacial advances throughout the region, some of which equaled or exceeded later LIA advances. The Altithermal interval, from ~ 11 ka to ~ 7 ka, was characterized by a lack of large glaciers. Neoglacial advances began as early as 8.4 ka with a series of progressively larger advances accompanied by retreating intervals. The Little Ice Age began to exert its influence on glacial advances as early as the 11th century AD, culminating by the 19th century in glacial extents greater than any since the latest Pleistocene (Menounos et al., 2009). The Canadian record is noteworthy for a remarkably robust data set. Glacial advances at Canadian latitudes often over-rode forests at low elevations, providing researchers with an accessible and radiocarbon-dateable record. One review study compiled 240 radiocarbon ages derived from exposed wood found in glacial forefields, in addition to radiocarbon dates from moraines and lake sediments (Menounos et al., 2009).

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However, a cautionary tale for over-interpretation of limited data is found at Mt. Waddington’s Tiedemann Glacier, where an early study documented a mid-Neoglacial expansion that greatly exceeded the glacier’s Little Ice Age moraines (Ryder and Thomson, 1986). The researchers consequently proposed a regional Tiedemann Advance. Subsequent studies throughout the Canadian Cordillera found evidence of glacial advances at similar times; however, no other glacier was found to exceed its LIA extent (Menounos et al., 2013). A recent investigation by Menounos et al. (2013) concluded that rock avalanche deposits significantly affected the behavior of the Tiedemann Glacier, enabling it to advance beyond its LIA extent. This example demonstrates the importance of finding a regionally relevant glacier to document significant advances and retreats.

SIERRA NEVADA AND SOUTH CASCADES In contrast to the extensive documentation in Canada, the research of Holocene glacial fluctuations in the southern ranges of the western North American Cordillera relies upon only a handful of studies. The glaciers of the Sierra Nevada advanced during the latest Pleistocene to their greatest post-LGM extent (Clark and Gillespie, 1997). Recent exposure dating of these moraines suggest retreat ages of 12.5 ka to 11.8 ka, demonstrating that the Holocene commenced with a warming climate (Marcott, 2011). Neoglacial advances began ~ 3.2 ka, with a series of pulses culminating in the Little Ice Age , which marked the maximum Holocene glacier extent (Bowerman and Clark, 2011). In the Oregon Cascades, where there is little documentation of Neoglacial activity, a study by Marcott et al. (2009) found evidence for a major advance prior to 8.0 ka, a poorly-constrained Neoglacial advance between 7.6 ka and 1.7 - 2.5 ka, and a LIA advance that peaked within the last 300 years. The early Neoglacial moraines were found ~ 100 m downslope of the LIA moraines, a rare instance of an earlier Neoglacial advance apparently exceeding the Little Ice Age advance.

NORTH CASCADES The Holocene glacial record in the North Cascades is more thorough than in Oregon but has been limited by a lack of dateable evidence. Traditional glacial fluctuation studies have focused on moraine dating (e.g., Davis et al., 2009), but this method only works where moraines are preserved. The advances of the Little Ice Age exceeded previous Neoglacial advances, and therefore have destroyed previous deposits. Shorn in-situ trees in the glacial forefield, which are so abundant in Canada, are extremely hard to find in the North Cascades, so radiocarbon dateable material is scarce. Recent studies have utilized cosmogenic radionuclide dating (Marcott, 2011), but this method is still limited by a lack of appropriate deposits, such as preserved moraines. In addition to limited data, the North Cascades record has considerable controversy regarding possible early Holocene advances. A succession of researchers have argued that glaciers at Mt. Baker advanced during the early Holocene (e.g., Thomas et al., 2000, Kovanen and Slaymaker, 2005). In response, a number of studies have been published that disagree with the timing of the purported early Holocene advances, instead placing these advances in the latest Pleistocene (e.g., Osborn et al., 2012). Other sites of contentious early Holocene advances include Glacier Peak (Beget, 1984, 1981), the Enchantment Lakes basin (Waitt Jr. et al., 1982), and Mt. Rainer (Heine, 1998). In opposition, Bilderback (2004) and Marcott (2011) placed the Glacier Peak / Enchantment

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Lakes advance in the latest Pleistocene, and Samolczyk (2011) did the same at Mt. Rainer. In their review of glacial and climate records from southern British Columbia and Washington, Osborn et al. (2012) concluded that the region did not see major early Holocene glacial advances, and furthermore that North Cascade glacial fluctuations follow similar trends throughout the region. The balance of evidence suggests a broadly coherent pattern of glacial fluctuations in the North Cascades, summarized by a large advance during the latest Pleistocene, retreat during the Altithermal, and periodic advances throughout the Neoglacial that culminated with the Little Ice Age advance (Figure 3). Questions include whether early Neoglacial advances inferred at Mt. Baker (Osborn et al., 2012), the Enchantment Lakes basin (Bilderback, 2004), and Mt. Rainer (Samolczyk, 2011) correlate throughout the region, as well as when the LIA advance began and when it reached its climax. Currently, much of the uncertainty in the North Cascades surrounding early Holocene and early Neoglacial advances is based upon interpretations of moraine records. Multi-proxy records would provide broader evidence for the timing of these proposed advances, and a high-resolution, continuous record would improve the fluctuation history considerably.

Glacial Fluctuation Proxy Records Holocene glacial researchers have used terrestrial and lacustrine proxy records to reconstruct fluctuation chronologies. Terrestrial proxy records typically focus on preserved glacial deposits (i.e., moraines) and dating of detrital and in-situ wood samples in glacial forefields (Davis et al., 2009; Menounos et al., 2009). Lacustrine proxy records are derived from glacigenic sediments preserved in proglacial lakes. Either method has its weaknesses and complications; used together, they can create a much stronger record (Osborn et al., 2007).

Terrestrial Records Well-preserved lateral and terminal moraines allow an accurate reconstruction of such crucial parameters as glacial extent, glacial volume and mass-balance (Osborn et al., 2007), as well as paleo-ELAs (Bowerman and Clark, 2011). They are especially useful because they provide relatively direct and accessible records of past fluctuations. In addition, moraine sets can be compared and relatively dated based on morphology, stratigraphic position, and vegetative cover of the moraines (Marcott et al., 2009). Moraines are dated using a variety of methods, including dendrochronology, lichenometry, tephrachronology, radiocarbon dating, and cosmogenic radionucleide (CRN) dating (Davis et al., 2009; Marcott, 2011; Marcott et al., 2009; Menounos et al., 2009; Osborn et al., 2007). Dating of moraines with dendrochronology or lichenometry is age-limited by tree and lichen lifespans, and in practice is most useful for Little Ice Age deposits. Well-dated tephras are helpful in constraining moraine ages, but are not always present. Radiocarbon dating of exposed organic remains is widely used in regions where glaciers have advanced below treeline, such as the western Canadian Cordillera, but has limited use elsewhere. The recent advent of high-precision CRN dating with 10Be holds great promise for dating Holocene moraines (Davis et al., 2009; Marcott, 2011).

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Dateable material found on top of a moraine (e.g., tephra, detrital wood, preserved organic debris, paleosols, or boulders for CRN) generally provide minimum ages of the advance that deposited the moraine, whereas material found under or within a moraine provide maximum ages (Davis et al., 2009; Marcott et al., 2009). In practice, this can lead to poorly constrained advance ages; Marcott et al. (2009) was only able to limit the age of a set of Neoglacial moraines to a six thousand year interval. Moreover, organic dateable material is often lacking, as alpine environments tend to have little organic activity and high energy levels. Relative dating methods allow moraine sets in different basins to be correlated, allowing regional glacial advances to be constructed. However, it must be noted that relying upon relative dating methods has given gross errors in chronologies, as evidenced by many early Holocene advances being re-dated as latest Pleistocene (Clark and Gillespie, 1997; Marcott, 2011). In addition to these dating limitations, the moraine record is hampered by its discontinuous nature. Lateral and terminal moraines do not record the duration of a fluctuation event; rather, they only record a “snapshot” of the glacier’s maximum extent, giving little evidence of how long the glacier may have occupied the position. Preservation problems also limit the moraine record’s continuity. Each glacial advance will remove any previous deposits it encounters, and the unconsolidated nature of glacial deposits renders them susceptible to rapid erosion. Thus, the moraine record is unavoidably biased towards the youngest and largest advances (i.e., Little Ice Age and Last Glacial Maximum), with little evidence of other advances. The terrestrial glacial fluctuation record also includes dateable material found in glacial forefields, such as in-situ glacially-killed trees, detrital wood, paleosols, tephras, and organic-rich sediments from bogs or ponds (Menounos et al., 2009). Glacial forefield records suffer from similar temporal limitations as moraine-derived records. In most cases, radiocarbon dates are considered maximum ages, as the glacier may have continued to advance beyond the sample location for hundreds to thousands of years. Sheared tree stumps in forefields are the strongest form of evidence, as they are easily dated and unequivocally record a glacier’s advance (Osborn et al., 2007). Detrital wood dating, in contrast, is complicated by uncertain provenance (Ryder and Thomson, 1986). For example, if a tree fell onto a retreating glacier and was delivered to the forefield, it might provide an age younger than the glacier’s advance age. Although useful, dateable forefields are limited temporally to glacial advances that sufficiently post-date the last advance for a new forest to grow, and they are limited spatially to locations below treeline. Most glacial sites in the coterminous United States have little to no dateable material in their forefields. The southern Coast Mountains of British Columbia, in contrast, have a particularly robust record of in-situ sheared stumps (Menounos et al., 2009).

Lacustrine Record In contrast to terrestrial records, proglacial lacustrine sediments can provide continuous, dateable and high temporal resolution records of glacier fluctuations. Karlén's (1976) work demonstrated the value of these records by noting that increasing glacial size correlates with increased downstream rock flour deposition in proglacial lakes. A succession of researchers have since refined the relationship between glacial activity and lacustrine sediments (Bakke et al., 2013, 2005; Hicks et al., 1990; Karlén, 1981, 1976;

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Leonard and Reasoner, 1999; Leonard, 1997, 1986; MacGregor et al., 2011; Maurer et al., 2012; Menounos et al., 2009; Osborn et al., 2007), with perhaps the strongest proof of concept coming from varved lake sediments in the Canadian Rockies (Leonard, 1997). Lacustrine sediments can be divided into three broad types; glacigenic clastic sediments (i.e., rock flour) that reflect upstream glacial erosion, non-glacigenic clastics from hillslope processes, and organic sediments. An ideal lacustrine setting would maximize the glacigenic signal, minimize the non-glacigenic clastic signal, and provide a steady- state background organic signal. Thus, the ratio of rock flour to organics becomes the proxy for glacier size. Crucially, a complete sedimentary record will reflect even minor glacial fluctuations for the duration of the record, and is therefore continuous (Leonard and Reasoner, 1999). Lacustrine sediments are typically dated by radiocarbon analysis of stratigraphically-preserved organic material. Varved sediments and tephra falls can be used for dating as well, but are not always present. Unlike many terrestrial environments, low-energy alpine lakes usually preserve dateable organic material (Leonard, 1986). Best practices rely upon dating terrestrial macrofossils (i.e., twigs, needles, leaves, seeds) instead of bulk organic sediments. Terrestrial macrofossils preserve original atmospheric ratios of radiocarbon, whereas bulk organic sediment samples can contain reservoir effects related to exogenous carbon (Leonard and Reasoner, 1999; Yu, 2014). The presence of carbonate bedrock or coal in a lake’s drainage basin would also hinder accurate dating by incorporating depleted carbon Interpretation of glacial lacustrine proxy records is complicated by sediment provenance and by the temporal relationship between glacial fluctuations and rock flour deposition. For example, in an extremely erodible environment, non-glacigenic clastic sediments were observed to overwhelm the rock flour signal (Hicks et al., 1990). Even in durable bedrock basins, clastic contributions from hillslope mass-wasting and other non- glacial processes must be assessed on a site-by-site basis. Organic sediment rates can trend over centuries in response to climate change, but do not fluctuate substantially over shorter time scales (Karlén, 1981; MacGregor et al., 2011). At a century and longer time scales, rock flour deposition reflects glacial activity (Leonard, 1997), but at finer temporal resolutions deposition can fluctuate in response to other variables (Leonard and Reasoner, 1999). At decadal time scales, there remains some question as to the relationship between rock flour deposition and glacial extent. Some studies indicate a correlation between the highest rates of deposition and transitional periods immediately before and after the maximum glacial extent (Leonard, 1997; Osborn et al., 2007). This may be a site-specific relationship; for example, a study by Bakke et al. (2005) of marine plateau glaciers observed a lag time of ten years between glacial expansion and rock flour deposition. In light of these complications, it is vital that the lacustrine sediment sources are well understood in the context of the lake’s drainage basin, and that interpretations on a sub-centennial timescale be well supported. The clarity of the rock flour signal can be enhanced or diminished by glacier and basin morphology. Glaciers that have melted completely can provide an unambiguous end-member example of a lacustrine sediment signal devoid of glacial inputs (Karlén, 1976). The benefit of such an “on/off” signal has been exploited where a glacier fluctuated across a hydrologic divide (Maurer et al., 2012), as well as where a glacier occupied a unique bedrock lithology (MacGregor et al., 2011). The position of the

7 Harold Wershow Thesis Proposal glacier in relation to the targeted lake also affects the rock flour signal. A small glacier’s rock flour signal can be diluted as it travels downstream by deposition in proglacial lakes, non-glacigenic inputs (fluvial and colluvial), and rock flour contributions from adjacent glaciers. A larger glacier, in contrast, may deliver too much sediment to proximal lakes, making it difficult to read minor fluctuations in the record. In a review of glacier and basin characteristics, Dahl et al. (2003) argued for a series of proglacial lakes that are fed by a single glacier, are retained by bedrock dams, and contain sites both proximal and distal to the glacier in order to accurately record signals at minimum and maximum glacial extents. The lacustrine record is strongest when it records deposition throughout the period of interest, is dateable by terrestrial macrofossils, tephras, and / or varves, and when the three respective sedimentary signals produce an unambiguous proxy of glacial fluctuations. However, site-specific complications can significantly impair the record. In comparison to the terrestrial record, it is a less direct proxy of glacial size and extent. Therefore, some researchers have argued for the benefit of collecting a multi-proxy record (Bakke et al., 2013; Osborn et al., 2007). In a review paper of glacial fluctuation studies in the Canadian Cordillera, Menounos et al. (2009) found broad agreement between terrestrial and lacustrine proxy records, noting that each record signaled a similar temporal onset of glacial advances. Ultimately, the lacustrine record is invaluable because it provides information concerning the onset, duration and magnitude of major and minor glacial fluctuations.

Equilibrium-Line Altitudes The goal of this project is to connect past glacial fluctuations to the climatic perturbations that drive them. Equilibrium-line altitudes (ELAs), which mark the elevation of the transition between a glacier’s accumulation and ablation zones, have been used by many researchers to link glacial fluctuations to climate change (Benn and Ballantyne, 2005; Bilderback, 2004; Bowerman and Clark, 2011; Leonard, 1989; Marcott et al., 2009; Sagredo et al., 2014). ELAs represent climatic conditions during a period of glacial equilibrium, when the glacier is neither advancing nor retreating. If the climate changes, then the ELA will raise or lower until the glacier is once again at equilibrium. Although ELAs are a widely-used glacial fluctuation proxy, glacial and climatic heterogeneities can complicate their interpretation. Large valley glaciers respond to regional climate fluctuations (Leonard, 1989), but they may not be as sensitive to smaller or shorter climate fluctuations. Conversely, smaller cirque glaciers tend to be more influenced by site-specific variables, such as aspect, concavity, and spatial relationship to peaks and passes (Graf, 1976). Changing climate conditions will have varying influence on ELA depending upon the regional climate (Sagredo et al., 2014). In wetter climates, ELAs responded dominantly to changes in temperature, whereas in drier climates, ELAs responded dominantly to changes in precipitation. An important implication of this study is that glaciers sharing a climatic regime will respond to regional climate change with similar ELA changes (Sagredo et al., 2014).

ELA RECONSTRUCTION Glacial fluctuation studies estimate ELAs of past glacier positions as a proxy for paleoclimate conditions. There are many methods for reconstructing paleo-ELAs, including

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Median Altitude, Lateral Moraine, Accumulation Area Ratio, and Toe-to-Headwall Altitude Ratio (Meierding, 1982), as well as Area x Altitude, Area x Altitude Balance Ratio, and Area x Altitude Balance Index (Osmaston, 2005). I will limit this discussion to the Lateral Moraine (LM), Accumulation Area Ratio (AAR), and Area x Altitude Balance Ratio (AABR) methods. The highest extent of lateral moraines (LM) is theoretically a direct indicator of paleo-ELA (Meierding, 1982). A glacier will only create lateral deposits below its ELA, in its ablation zone, because that is where ice is melting and therefore debris is being shed. However, because deposits may not be completely preserved, the maximum elevation of the remnant lateral moraine may be lower than the actual paleo-ELA (Bowerman and Clark, 2011). The LM method is further limited because short-lived glacier advances may not have sufficient time to deposit lateral moraines, leaving no record behind. Reconstruction of multiple paleo-ELAs is difficult because distinct multiple sets of lateral moraines are rarely preserved on steep valley walls. In contrast, terminal and recessional moraines are more frequently preserved as distinct sets because they are usually found on valley floors, a more stable environment. Preservation of distinct moraine sets allows reconstruction of multiple previous glacial extents. The Accumulation Area Ratio (AAR) method, which assumes a consistent ratio between the area of the accumulation zone and the ablation zone area, is commonly used in these scenarios because it only requires that a glacier’s outline be accurately mapped (Osmaston, 2005). The ratio used is regionally-specific; previous workers near the North Cascades have used an AAR of 0.65 (Bilderback, 2004; Marcott et al., 2009). However, this method assumes that a glacier’s net mass-balance, which is the difference between annual accumulation and ablation at a specific point, is invariant over the span of the glacier. For most glaciers, net mass-balance increases with distance from the ELA, meaning that a glacier’s terminus and head contribute more to its zones of ablation and accumulation, respectively, than equivalent areas closer to the ELA (Benn and Ballantyne, 2005). Therefore, the AAR method over-counts the area near its ELA, and under-counts the areas farthest from the ELA. A glacier’s unique hypsometry (area-elevation relationship) can also influence its mass-balance, an effect often seen in small alpine glaciers (Graf, 1976; Leonard, 1989; Osmaston, 2005). To explicitly account for glacial hypsometry as well as varying net mass-balance, Osmaston (2005) developed the Area x Altitude Balance Ratio (AABR) method. Using reconstructed surface contours, the method weights the mean altitude of each contour interval by the glacial surface area; thus, the Area x Altitude. The Balance Ratio is defined as the ratio between the mass-balance gradient (mass balance / altitude) above the ELA over the mass-balance gradient below the ELA. The AABR can readily be calculated from an Excel spreadsheet using a publicly available program (Osmaston, 2005). The choice of an ELA reconstruction method depends on site-specific conditions, most crucially preservation of moraines and glacial hypsometry (Meierding, 1982; Osmaston, 2005). Virtually all methods rely upon a reconstruction of a glacier’s paleo- extent (i.e., the area contained within the headwall, lateral moraines, and terminal moraine). As with fluctuation proxy records, multiple methods provide stronger support for interpretations.

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Objectives The timing and magnitude of regional glacial advances and retreats (fluctuations) in the North Cascades remains poorly constrained and controversial. The Lyman Glacier is particularly well suited to refining this uncertainty because it has promising terrestrial and lacustrine proxy records as well as modern glacial and climate records. The overall goals of my project are to establish a representative, continuous and well-dated record of Holocene glacial fluctuations for the North Cascades using the Lyman Glacier, and to use that record as a paleoclimate proxy by calibrating it against modern observations, and to synthesize Holocene and modern records to aid projections of glacial resources for the region.

Geologic Setting – Lyman Glacier & Basin

Basin Morphology The Lyman Glacier occupies the upper (southern) portion of a two-tiered north-northwest oriented basin (informally named the Lyman basin) that is ~ 5 km long and up to 2 km wide (Figure 1). The floor of the upper basin is ~ 1810 m high, whereas the floor of the lower (northern) portion, which is occupied by Lyman Lake, is ~ 1710 m high. The lake’s outflow becomes Railroad Creek and drains eastward into Lake Chelan after ~25 km and a drop in elevation of ~ 1325 m. The basin is defined by a high (~ 2400 m) ridgeline on the southern and western margins, forming an “L” that shelters the glacier from southwest exposure (Figure 4). The ridgeline is interrupted by Spider Gap ( ~ 300 m below ridgeline) to the southwest, and resumes at a lower elevation (~ 2000 m) on the east side of the basin. The glacier itself was measured at ~ 0.20 km2 in 2008 with an estimated volume of 9.8 million m3 (Pelto, 2009). It flows into the basin from a cirque (elev. ~1800 - 2000 m) on the northern flank of Chiwawa Mountain (Figure 5) and is currently calving into a shallow glacial lake. Above the glacier to the southwest lies a northeast-oriented snowfield that extends nearly to the bounding ridgeline of Chiwawa Mountain. Other relevant snowfields include the Spider Gap snowfield to the southeast as well as numerous smaller snowfields that line the high western ridgeline of the basin (Figure 4). The areas occupied by these snowfields may contribute snow via avalanches to the glacier below. The upper portion of the basin was recently occupied by a much larger glacier. Observable evidence (Figure 5) includes fluted till in the glacial forefield and a well- defined, un-vegetated terminal moraine set that juxtaposes mature alpine vegetation to the north (downstream) with immature vegetation to the south (Jumpponen et al., 1998). The moraines are most likely from the Little Ice Age advance; a photograph from 1901 by W.D. Lyman shows the glacial terminus at, or at least very near to, the terminal moraine (Figure 6). The lateral extent of the presumed Little Ice Age glacier is evidenced by the presence of orange (iron-oxidized) rocks high on the otherwise gray eastern scree slopes (Figure 5), as well as a subtle trimline on the western scree slopes. The glacial forefield contains a series of shallow (< 2 m deep) pools that channel glacial outflow from the calving terminus towards Lyman Lake.

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The lower portion of the basin surrounds Lyman Lake (elev. 1710 m), which is fed by meltwater from Lyman Glacier. There are no other significant fluvial contributions to the lake. The lake appears to be shallow, as evidenced by a likely moraine submerged less than two meters underwater (pers. comm., Ken Dull, USFS, 2014) that divides the lake into two sub-basins (Figure 7). The upper sub-basin (~ 0.16 km2) contains a prograding delta and is at the base of an extensive talus field on the northeastern flank of Chiwawa Mountain’s northern ridge. The lower sub-basin (~ 0.12 km2) is surrounded on three sides by forests on gentler slopes, with no evident clastic debris sources. However, there is a prominent southwest-facing cirque (elev. ~ 1900 – 2000 m) north of the lake that may have harbored a pocket glacier in the past (Figure 4).

Bedrock Geology The North Cascades bedrock geology is characterized by pre-Cambrian to upper Paleozoic highly metamorphosed rocks intruded by plutons from the early Triassic to Miocene (Cater, 1982). The Lyman basin is surrounded by a labradorite granodiorite (grannogabro) from the Tertiary (20-23 Ma) Cloudy Pass pluton. The grannogabro has mafic compositions ranging from ~ 15% - ~30%, can grade into labradorite quartz diorite or labradorite quartz monzonite, is often pseudoporphyritic with large labrodorite and quartz crystals, and is structurally characterized by vertical to near-vertical joint sets (Cater and Crowder, 1967). Chiwawa Mountain, however, is composed of the early Paleozoic Swakane Biotite Gneiss, and Spider Gap exposes a hornblende schist and gneiss that structurally overlies the Swakane (Figure 8). The Swakane has swirled gneissic foliation as well as abundant intrusions of leucratic quartz diorite. These sills, dikes, and irregular masses are commonly observed at both Chiwawa Mountain and Spider Gap. The metamorphosed rocks of Chiwawa Mountain and Spider Gap appear to be the source of the iron-oxidized orange clasts that define the trimline on the basin’s bounding eastern ridge (Figure 5).

Glacial & Instrumental Record The existing glacial record is remarkably detailed, yet only extends back to the late 1890s. Historical documentation begins with photographs of the glacier, lake and basin. Chelan County Public Utility District’s predecessor began collecting annual snow pack measurements (snow course) in 1926, a project continued to 1979, at which point a remotely operated snow telemetry site was installed (Figure 9), capable of measuring snow pack as well as temperature (Buehn, 2014). This site is fortuitously located at the same elevation (~ 1800 m) as the glacier and is only 3.5 km away. Beginning in 1929 and continuing through at least 1940, the Washington Water Power Company visited the glacier annually to record the terminus position (Freeman, 1941). Photographic documentation continued throughout the 20th century, including aerial photography by the U.S. Geological Survey. The first prominent scientific report concerning the Lyman Glacier came in 1941, when Freeman reported on the glacier’s retreat using the Washington Water Power Company’s records. Direct measurements of the glacier’s terminus were made intermittently for the remainder of the 20th century (Pelto and Hedlund, 2001). Beginning in 1986, direct mass-balance measurements have been taken annually at Lyman Glacier, a

11 Harold Wershow Thesis Proposal project that continues to the present (Pelto and Riedel, 2001). The glacier’s 20th century history has been reconstructed by Jumpponen et al. (1998), who noted the conspicuous lack of recessional moraines in the glacial forefield, as well as by Pelto (2009). In contrast, Freeman in 1941 reported observing numerous recessional moraines. It appears that the early 20th century recessional moraines have deteriorated considerably. The glacier’s retreat from its Little Ice Age maximum extent was first observed in 1898 (Jumpponen et al., 1998). A 1901 photograph illustrates that the glacier was quite close its terminal moraine (Figure 6). The three-dimensional extent of the glacier is well- displayed by a 1915 photograph, which shows the glacier filling the entire valley (Figure 10). A 1921 photograph illustrates the continuity between the main body of the glacier and outlying snowfields below Spider Gap and Chiwawa Mountain (Figure 11). In contrast, a 1997 photograph highlights the isolation of the modern glacier from these snow sources (Figure 12). Jumpponen et al.'s (1998) reconstruction outlines the glacier’s two-dimensional extent, whereas Pelto’s (2009) reconstruction highlights reductions in areal extent, volume and terminus position (Figure 13, A-C). The Lyman Glacier has steadily retreated throughout the 20th century, even between 1950 and 1976, a period that saw many North Cascades glaciers advancing (Pelto and Hedlund, 2001). At its present rate of retreat, it is expected to disappear before 2040 (Pelto, 2009).

Methods

Geomorphic and Geologic Mapping Past geomorphic mapping of the Lyman basin has focused on the glacial forefield, leaving the rest of the basin unmapped. Freeman (1941) mapped the glacier’s retreat, noting the constantly changing location of the terminus as well as short-lived recessional moraines. More recently, the glacier’s retreat has been mapped by Jumpponen et al. (1998) and Pelto (2009). However, there is no indication that any workers have examined the basin north of the Little Ice Age terminal moraine. Previous workers have mapped the local bedrock geology at a 1:62,500 scale (Cater and Crowder, 1967); however, this map has not been updated since 1967 and is insufficiently detailed for my purposes. I will map the geomorphic evidence of past glacial activity in order to reconstruct past glacial extents. I will map bedrock geology as well, to identify possible sediment sources for the lacustrine record.

GEOMORPHIC MAPPING Initial reconnaissance mapping with aerial photographs will prioritize locations to be mapped and sampled in the field. I will identify surficial structures such as moraines, outwash channels, and hillslope deposits using low snow-cover stereoscopic images from 1984 via the USGS’s Aerial Photo Single Frame project and from 1990 via the USGS’s National Aeronautic Photography Program. Additionally, an orthorectified high resolution (0.5 m) color set of images from the USGS OrthoImagery collection will be used to assist with GIS mapping. Using a 10m DEM overlain by the existing 7.5’ geologic map (Cater and Crowder, 1967), I will create a field map that highlights the relevant geologic and geomorphic features of the area.

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Geomorphic field mapping will provide additional constraints on the reconnaissance mapping to improve my interpretations of past glacial extent. Besides verifying feature types (e.g., confirming nature and continuity of moraines), I will evaluate relative weathering of glacial deposits (e.g., Bowerman and Clark, 2011; Davies, 2011). By combining my relative age interpretations with observed stratigraphic positions, I will attempt to create a relative chronology of prominent glacial fluctuations in the Lyman basin. I will also search the upper portions of the basin for past sources of snow and rock fall into the basin. Previous extents of the Lyman Glacier had multiple snow sources (Figures 6, 10, 11). For example, a 1915 photograph shows significant flow onto the glacier from the Spider Gap glacier (Figure 10), which later became disconnected from Lyman and is now a snowfield (Figure 12). I will investigate possible source areas on either side of the basin, looking for evidence of past glacial activity such as striated bedrock and moraines. I will also assess hillslope rock deposits in the upper basin that may have influenced the glacier (e.g., Menounos et al., 2013) as well as deposits in the lower basin that may have impacted the lake.

GEOLOGIC MAPPING Although the existing geologic map gives gross unit delineation, spatially precise bedrock identification is crucial for determining the possible provenance of lake sediments. Depending on the rock type, the lacustrine rock flour will have a traceable mineralogical composition. For example, MacGregor et al. (2011) used the intermittent presence of carbonates in their sediment core stratigraphy to infer glacial fluctuations that advanced over a discrete patch of dolomite bedrock in the forefield. The presence of three distinct rock types near the Lyman Glacier’s accumulation zone (Figure 8) may present a similar opportunity. As evidenced by the oxidized eastern trimline (Figure 5), the metamorphic rocks can provide a distinct visual signal. Furthermore, rock flour parameters such as Magnetic Susceptibility will be influenced by the composition of the bedrock, making detailed knowledge of the source composition vital to interpreting the rock flour signal. In addition to field checking Cater and Crowder’s (1967) geologic map, I will locate the transitions between rock units in the upper basin headwall and assess the circumstances under which each unit may have delivered sediments to the downstream lakes.

Lacustrine Glacial Fluctuation Record

LACUSTRINE SEDIMENT CORING The lacustrine record complements the terrestrial record by providing a continuous signal of relative glacial extent throughout the Holocene. The lacustrine record tends to be best- preserved and extend for the longest duration in the deepest area of a lake (e.g., Davies, 2011). Because the lake depths are unknown, I will collect bathymetric data from Lyman Lake as well as the upstream proglacial lakes using an inflatable raft, a differential GPS, and a handheld bathymetry meter. I will create detailed bathymetric maps using Surfer (v. 8.0), in order to target the most promising coring sites (e.g., Bowerman and Clark, 2011).

13 Harold Wershow Thesis Proposal

I will use two types of coring systems: a Glew corer (Glew, 1991, 1988) and a Livingstone piston corer (Wright, 1967). The Glew corer provides high-resolution sampling of young near-surface sediments, which tend to be poorly-consolidated, whereas the Livingstone corer is designed to recover deeper sediments that might span the duration of the Holocene. For both systems, coring will take place from a floating platform anchored above the target site. Adjacent lakes and bogs that are not fed by rock flour will be cored as well to provide complementary records of non-glacial sediments (Dahl et al., 2003).

SEDIMENTOLOGY My analyses of the sediment cores will focus on discriminating and quantifying glacial versus. non-glacial sediment fluxes to the lakes. My sediment analysis will begin at LacCore (University of Minnesota), following their Initial Core Description process, which includes visual core stratigraphy description, magnetic susceptibility, gamma density, and loss-on-ignition. Visual stratigraphy will highlight macro-scale changes in sediment composition as well as quantifying sediment composition via smear slide analysis (pers. comm.., Anders Noren, LacCore). Magnetic susceptibility can be a proxy for glacigenic mineral presence, depending on the composition of the source rock (Bakke et al., 2013). Gamma density allows for high-resolution density measurements that correlate with clastic content (pers. comm.., Anders Noren, LacCore ). Loss-on-ignition (LOI) is a direct measurement of the proportion of organic versus minerogenic sediment in a sample (Bakke et al., 2013). Non-glacigenic clastic sediments are expected to have much larger grain-size than rock flour, allowing grain-size analysis to discriminate the sources of minerogenic sediments (MacGregor et al., 2011). Grain-size analysis will take place at WWU, using the Malvern Mastersizer 2000. Each parameter is potentially useful as a direct proxy for glacial fluctuations (Bakke et al., 2005), but multiple interpretations are possible. Bakke et al. (2013) found strong statistical correlations among these parameters and advocate for multiple parameter analysis. I will compare the results from each analysis and synthesize them to construct a coherent record of glacial fluctuations.

SEDIMENT DATING The glacial fluctuation proxy record developed via sediment analysis must be temporally linked to the core stratigraphy. I will develop a stratigraphic age model using 210Pb analysis, tephrochronology, and accelerator mass spectrometry radiocarbon analysis. 210Pb analysis is useful for high-resolution dating of young (< 200 years) sediments (Appleby and Oldfield, 1978). This will allow the recent sedimentary record to be precisely matched to the existing high-resolution historical record. Radiocarbon analysis has a coarser age resolution but will provide age constraints for the remainder of the sediment core. I will search for terrestrial macrofossils in preference to bulk sediments to avoid radiocarbon reservoir effects (Yu, 2014). Terrestrial macrofossils will be separated from cores and dried to preserve their integrity shortly after initial core analysis (Wohlfarth et al., 1998). I expect to find several tephra layers, including ash falls from Mt. Mazama, Glacier Peak and Mt. St. Helens. Positive identification of well-dated tephra falls will improve the accuracy of the stratigraphic age model. 210Pb analysis and tephrochronology will be done at WWU. I will travel to Lawrence Livermore National

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Laboratory (LLNL) to prepare my radiocarbon samples. The combined results of these analyses will be used to construct a stratigraphic age model, thereby dating the glacial fluctuation record (e.g., Bowerman and Clark, 2011; Davies, 2011).

Calibration of Glacial Record to Climate Record The Lyman Glacier’s dated fluctuation record will provide a proxy record for paleoclimate by using equilibrium-line altitudes. Historical (post-1890s) ELAs have been directly measured since 1986 (Pelto and Riedel, 2001) and can readily be reconstructed since the Little Ice Age using photographs and terminus observations. I anticipate using the Lateral Moraine, Accumulation Area Ratio (AAR), and Area x Altitude Balance Ratio (AABR) methods. I will use an AAR of 0.65 (Bilderback, 2004; Marcott et al., 2009), and balance ratios from 1 to 3 for the AABR method (Marcott et al., 2009). The historical climate records (SNOTEL data since 1979 and snow course data since 1926) will be linked to historical ELAs, allowing an empirical relationship to be developed between three crucial parameters: Mean Summer Temperature, Cumulative Winter Precipitation, and ELA. Because there will be a complementary high-resolution glacial fluctuation proxy record (i.e., rock flour), the Lyman Glacier’s historical ELA fluctuations can be empirically linked to its fluctuation record. By combining these two relationships, climate parameters can be empirically tied to rock flour fluctuations, and thus the modern climate record will be calibrated to the modern glacial record. The derived empirical relationships between rock flour and climate will be applied to the entire rock flour record, creating a complementary climate record for the duration of the Holocene. This methodology is uniquely suited to the Lyman Glacier because of the remarkable historical glacial and climate records. However, because it is novel, I will check it with an established method. Following the methods of Bilderback (2004), Bowerman and Clark (2011), and Marcott et al. (2009), I will plot climate conditions for measured and reconstructed ELAs on Leonard's (1989) “envelope”, which is an empirical plot of climate conditions at the ELAs of 32 modern glaciers (Figure 14). The temperature / precipitation values necessary for a glacier to exist, derived from its modern ELA, are compared to the present-day temperature / precipitation conditions at a reconstructed paleo-ELA. The climatic difference between the two ELAs is equivalent to the change that has occurred in the glacier’s climate, and can be easily calculated graphically. Thus, for every reconstructed paleo-ELA, an additional climate data point can be added to the glacier’s climate record. Besides the reconstructed ELAs from the LIA maximum extent to present, additional paleo-ELAs will be developed from the lacustrine record. Minimums in rock flour deposition reflect time periods when the glacier had entirely melted. The paleo- ELAs must have been greater than 2000 meters high at these times. Thus, rock flour minimums will equate with modern temperature / precipitation conditions at > 2000 m of elevation. The additional data points on the temperature / precipitation “envelope” will help constrain the Lyman Glacier’s projected climate record.

Anticipated Results I expect the lacustrine record to reflect known climate trends of the Holocene. Major trends include early Holocene warming, mid- to late-Holocene cooling, maximum

15 Harold Wershow Thesis Proposal cooling during the Little Ice Age, and abrupt warming in the 20th century. The rock flour signal should be minimal during the early Holocene Altithermal period, increase during the mid- to late-Holocene Neoglacial period, peak during the Little Ice Age, and wane dramatically in the 20th century. On a finer scale, I expect to find pulses in the rock flour signal that correlate to known Neoglacial events. These might include the initial Neoglacial advance at 6.0 ka observed at the Easton Glacier on Mt. Baker by Osborn et al. (2012), the 3.3 - 2.8 ka advance noted by Bilderback (2004) in the nearby Enchantments Lakes basin, and the four late Holocene advances observed at the Deming Glacier on Mt. Baker (Osborn et al., 2012). A peak rock flour signal, reflective of maximum glacial extent, would be evidenced by a high ratio of glacial sediments to non-glacial sediments. I would expect to see corresponding peaks in magnetic susceptibility and gamma density and a trough in loss-on-ignition measurements, whereas grain-size analysis would probably show a peak in clay and silt composition. I would expect these signals to be inverted during minimum glacial extents.

Potential Problems Lake sediment coring always presents challenges, particularly in remote settings. Physically transporting the required equipment to and from the field site will be arduous, but can be accomplished with multiple trips. I am also investigating the feasibility of hiring pack horses. If I cannot afford them, I will rely upon the broad backs and strong legs of many willing field assistants. I have nine field assistants committed for various durations thus far, and I hope to have enlisted at least three more by the end of the summer. Lake sediment coring is logistically complicated and requires mild weather for success. Accordingly, coring will occur in mid-August and I will have the assistance of three experienced corers. If we are not successful, I have planned a second coring campaign for early September. I expect that interpreting the rock flour signal will be the greatest challenge of this project. Although the sediment parameters are well-understood (e.g., high LOI equates with minimal glacial presence), every site presents its own complications. For example, if the ratio of organic to clastic sediments becomes too low (<5% LOI), LOI measurements lose their sensitivity to glacial fluctuations (Bakke et al., 2005). The Mazama tephra, which I expect find in Lyman Lake, has been observed to artificially increase the clastic sediment ratio due to post-deposition reworking (Osborn et al., 2007). The rock flour signal is known to lag behind glacial pulses, but the timing appears to vary from site to site (Bakke et al., 2005; Leonard, 1997; Menounos et al., 2009; Osborn et al., 2007). However, the Lyman Glacier’s record is fundamentally robust because these site- specific variables can be controlled by instrumental records. Climate conditions, ELAs, and the rock flour signal will all be correlated on a decadal scale for the last 90+ years, a period that encompasses nearly the entire range of Holocene variability in glacial extent.

Work and Research Plan Field work will commence, and hopefully conclude, in the summer of 2014. I will map the bathymetry from July 4th to July 6th, I will map the geomorphology and bedrock

16 Harold Wershow Thesis Proposal geology from August 7th to August 15th, and I will collect sediment cores from August 18th to August 22nd. I will return in the last week of August to address any outstanding tasks. If necessary, I will return once again during the second week of September to collect more sediment cores. Laboratory analysis will begin at LacCore during the 3rd week of September, and is expected to take a full week. In addition to the Initial Core Description process, I will select and preserve samples for radiocarbon dating. Upon return to Western Washington University, I will conduct 210Pb analysis, grain-size analysis and tephrochronology, although the timing is flexible. Radiocarbon analysis will occur according to the schedule of LLNL, presumably at some point in the fall or winter of 2014. Interpretation of results will take place in 2015.

References Appleby, P.G., Oldfield, F., 1978. The calculation of lead-210 dates assuming a constant rate of supply of unsupported 210Pb to the sediment. CATENA 5, 1–8. doi:10.1016/S0341-8162(78)80002-2 Bakke, J., Lie, Ø., Nesje, A., Dahl, S.O., Paasche, Ø., 2005. Utilizing physical sediment variability in glacier-fed lakes for continuous glacier reconstructions during the Holocene, northern Folgefonna, western Norway. The Holocene 15, 161–176. doi:10.1191/0959683605hl797rp Bakke, J., Trachsel, M., Kvisvik, B.C., Nesje, A., Lyså, A., 2013. Numerical analyses of a multi-proxy data set from a distal glacier-fed lake, Sørsendalsvatn, western Norway. Quat. Sci. Rev. 73, 182–195. doi:10.1016/j.quascirev.2013.05.003 Beget, J.E., 1981. Early Holocene glacier advance in the North Cascade Range, Washington. Geology 9, 409–413. doi:10.1130/0091- 7613(1981)9<409:EHGAIT>2.0.CO;2 Beget, J.E., 1984. Tephrochronology of late Wisconsin deglaciation and Holocene glacier fluctuations near Glacier Peak, North Cascade Range, Washington. Quat. Res. 21, 304–316. doi:10.1016/0033-5894(84)90070-X Benn, D.I., Ballantyne, C.K., 2005. Palaeoclimatic reconstruction from Loch Lomond Readvance glaciers in the West Drumochter Hills, Scotland. J. Quat. Sci. 20, 577– 592. doi:10.1002/jqs.925 Bilderback, E., 2004. TIMING AND PALEOCLIMATIC SIGNIFICANCE OF LATEST PLEISTOCENE AND HOLOCENE CIRQUE GLACIATION IN THE ENCHANTMENT LAKES BASIN, NORTH CASCADES, WA (Master). Western Washington University, Bellingham, WA. Bowerman, N.D., Clark, D.H., 2011. Holocene glaciation of the central Sierra Nevada, California. Quat. Sci. Rev. 30, 1067–1085. doi:10.1016/j.quascirev.2010.10.014 Buehn, S.J., 2014. Snow Survey Data. Cater, F.W., 1982. Intrusive rocks of the Holden and Lucerne quadrangles, Washington; the relation of depth zones, composition, textures, and emplacement of plutons (No. PP - 1220). United States Geological Survey. Cater, F.W., Crowder, D.F., 1967. Geologic map of the Holden quadrangle, Snohomish and Chelan Counties, Washington. Geologic Quadrangle Map GQ-646.

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Clark, D.H., Gillespie, A.R., 1997. Timing and significance of Late-glacial and Holocene cirque glaciation in the Sierra Nevada, California. Quat. Int. 38–39, 21–38. doi:10.1016/S1040-6182(96)00024-9 Clark, P.U., Dyke, A.S., Shakun, J.D., Carlson, A.E., Clark, J., Wohlfarth, B., Mitrovica, J.X., Hostetler, S.W., McCabe, A.M., 2009. The Last Glacial Maximum. Science 325, 710–714. doi:10.1126/science.1172873 Dahl, S.O., Bakke, J., Lie, Ø., Nesje, A., 2003. Reconstruction of former glacier equilibrium-line altitudes based on proglacial sites: an evaluation of approaches and selection of sites. Quat. Sci. Rev. 22, 275–287. doi:10.1016/S0277- 3791(02)00135-X Davies, N., 2011. Holocene glaciation of the Green River drainage, Wind River Range, Wyoming (M.S.). Western Washington University, Bellingham, WA. Davis, P.T., Menounos, B., Osborn, G., 2009. Holocene and latest Pleistocene alpine glacier fluctuations: a global perspective. Quat. Sci. Rev. 28, 2021–2033. doi:10.1016/j.quascirev.2009.05.020 Freeman, O.W., 1941. The Recession of Lyman Glacier Washington. J. Geol. 49, 764– 771. Glew, J.R., 1988. A portable extruding device for close interval sectioning of unconsolidated core samples. J. Paleolimnol. 1, 235–239. doi:10.1007/BF00177769 Glew, J.R., 1991. Miniature gravity corer for recovering short sediment cores. J. Paleolimnol. 5, 285–287. doi:10.1007/BF00200351 Graf, W.L., 1976. Cirques as Glacier Locations. Arct. Alp. Res. 8, 79–90. doi:10.2307/1550611 Heine, J.T., 1998. Extent, timing, and climatic implications of glacier advances Mount Rainer, Washington, U.S.A., at the Pleistocene/Holocene Transition. Quat. Sci. Rev. 17, 1139–1148. doi:10.1016/S0277-3791(97)00077-2 Hicks, D.M., McSaveney, M.J., Chinn, T.J.H., 1990. Sedimentation in Proglacial Ivory Lake, Southern Alps, New Zealand. Arct. Alp. Res. 22, 26–42. doi:10.2307/1551718 Jumpponen, A., Mattson, K., Trappe, J.M., Ohtonen, R., 1998. Effects of Established Willows on Primary Succession on Lyman Glacier Forefront, North Cascade Range, Washington, U.S.A.: Evidence for Simultaneous Canopy Inhibition and Soil Facilitation. Arct. Alp. Res. 30, 31–39. doi:10.2307/1551743 Karlén, W., 1976. Lacustrine Sediments and Tree-Limit Variations as Indicators of Holocene Climatic Fluctuations in Lappland, Northern Sweden. Geogr. Ann. Ser. Phys. Geogr. 58, 1. doi:10.2307/520740 Karlén, W., 1981. Lacustrine Sediment Studies. A Technique to Obtain a Continous Record of Holocene Glacier Variations. Geogr. Ann. Ser. Phys. Geogr. 63, 273. doi:10.2307/520840 Kovanen, D.J., Easterbrook, D.J., 2001. Late Pleistocene, post-Vashon, alpine glaciation of the Nooksack drainage, North Cascades, Washington. Geol. Soc. Am. Bull. 113, 274–288. doi:10.1130/0016-7606(2001)113<0274:LPPVAG>2.0.CO;2 Kovanen, D.J., Slaymaker, O., 2005. Fluctuations of the Deming Glacier and theoretical equilibrium line altitudes during the Late Pleistocene and Early Holocene on

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Mount Baker, Washington, USA. Boreas 34, 157–175. doi:10.1111/j.1502- 3885.2005.tb01012.x Leonard, E.M., 1986. Use of lacustrine sedimentary sequences as indicators of Holocene glacial history, Banff National Park, Alberta, Canada. Quat. Res. 26, 218–231. doi:10.1016/0033-5894(86)90106-7 Leonard, E.M., 1989. Climatic Change in the Colorado Rocky Mountains: Estimates Based on Modern Climate at Late Pleistocene Equilibrium Lines. Arct. Alp. Res. 21, 245–255. doi:10.2307/1551563 Leonard, E.M., 1997. The relationship between glacial activity and sediment production: evidence from a 4450-year varve record of neoglacial sedimentation in Hector Lake, Alberta, Canada. J. Paleolimnol. 17, 319–330. doi:10.1023/A:1007948327654 Leonard, E.M., Reasoner, M.A., 1999. A Continuous Holocene Glacial Record Inferred from Proglacial Lake Sediments in Banff National Park, Alberta, Canada. Quat. Res. 51, 1–13. doi:10.1006/qres.1998.2009 Lyman, W.D., 1909. The Columbia River Its History, Its Myths, Its Scenery, Its Commerce, Project Gutenberg EBook. G.P. Putnam’s Sons; The Knickerbocker Press, New York. MacGregor, K.R., Riihimaki, C.A., Myrbo, A., Shapley, M.D., Jankowski, K., 2011. Geomorphic and climatic change over the past 12,900 yr at Swiftcurrent Lake, Glacier National Park, Montana, USA. Quat. Res. 75, 80–90. doi:10.1016/j.yqres.2010.08.005 Marcott, S.A., 2011. Late Pleistocene and Holocene Glacier and Climate Change (Ph.D.). Oregon State University, United States -- Oregon. Marcott, S.A., Fountain, A.G., O’Connor, J.E., Sniffen, P.J., Dethier, D.P., 2009. A latest Pleistocene and Holocene glacial history and paleoclimate reconstruction at Three Sisters and Broken Top Volcanoes, Oregon, U.S.A. Quat. Res. 71, 181–189. doi:10.1016/j.yqres.2008.09.002 Marcott, S.A., Shakun, J.D., Clark, P.U., Mix, A.C., 2013. A Reconstruction of Regional and Global Temperature for the Past 11,300 Years. Science 339, 1198–1201. doi:10.1126/science.1228026 Maurer, M.K., Menounos, B., Luckman, B.H., Osborn, G., Clague, J.J., Beedle, M.J., Smith, R., Atkinson, N., 2012. Late Holocene glacier expansion in the Cariboo and northern Rocky Mountains, British Columbia, Canada. Quat. Sci. Rev. 51, 71–80. doi:10.1016/j.quascirev.2012.07.023 Meierding, T.C., 1982. Late pleistocene glacial equilibrium-line altitudes in the Colorado Front Range: A comparison of methods. Quat. Res. 18, 289–310. doi:10.1016/0033-5894(82)90076-X Menounos, B., Clague, J.J., Clarke, G.K.C., Marcott, S.A., Osborn, G., Clark, P.U., Tennant, C., Novak, A.M., 2013. Did rock avalanche deposits modulate the late Holocene advance of Tiedemann Glacier, southern Coast Mountains, British Columbia, Canada? Earth Planet. Sci. Lett. 384, 154–164. doi:10.1016/j.epsl.2013.10.008 Menounos, B., Osborn, G., Clague, J.J., Luckman, B.H., 2009. Latest Pleistocene and Holocene glacier fluctuations in western Canada. Quat. Sci. Rev. 28, 2049–2074. doi:10.1016/j.quascirev.2008.10.018

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Figures

FIGURE 1 – GEOGRAPHICAL SETTING OF THE LYMAN GLACIER

FIGURE 2 – WESTERN NORTH AMERICAN CORDILLERA

FIGURE 3 – SUMMARIZED GLACIAL FLUCTUATION HISTORIES

FIGURE 4 – BASIN MORPHOLOGY OF THE LYMAN GLACIER

FIGURE 5 – GLACIAL FOREFIELD

FIGURE 6 – PHOTOGRAPH OF LYMAN GLACIER IN 1901

FIGURE 7 – LYMAN LAKE

FIGURE 8 – BEDROCK GEOLOGY OF THE LYMAN GLACIER BASIN

FIGURE 9 - LYMAN LAKE SNOTEL SITE

FIGURE 10 – PHOTOGRAPH OF LYMAN GLACIER IN 1915

FIGURE 11 – PHOTOGRAPH OF LYMAN GLACIER IN 1921

FIGURE 12 – PHOTOGRAPH OF LYMAN GLACIER IN 1997

FIGURE 13 (A-C) – RECESSION OF LYMAN GLACIER DURING THE 20TH CENTURY

FIGURE 14 – CLIMATE CONDITIONS FOR MODERN GLACIERS

21 Lake Chelan 1 km Glacier Peak N

Washington Lyman basin

Lyman Lake

Lyman Glacier

FIGURE 1 – GEOGRAPHICAL SETTING OF THE LYMAN GLACIER Lyman Glacier is located in the Glacier Peak Wilderness area of the Wenatchee-Okanogan National Forest. It is 17 km northeast of Glacier Peak. Lyman Lake, 3 km downstream of the glacier, drains into Lake Chelan via the Railroad Creek drainage, a path of ~ 25 km. Western Canadian 2 Cordillera (1) 3 4 5 North Cascades 6 7

South Cascades 8

Sierra Nevada (10) 9

FIGURE 2 –WESTERN NORTH AMERICAN CORDILLERA Numbered blue stars align with locations in Figure 3, listed from North to South. The emphasis on the “western” North American Cordillera is to distinguish between glaciers in uenced by maritime moisture and those aected by dry, continental conditions. My study will focus on the former. For a discussion of continental glaciers of the North American Cordillera, see Marcott, 2011; Maurer et al., 2012; and Menounos et al., 2009. 11.7ka 9ka 6ka 3ka 1ka 0ka

1: Western Canadian Cordillera, B.C. (Menounos, et al., 2009) 2: Tiedemann Glacier, Mt. Wadding- ton, B.C. (Menounos, et al., 2013) 3: Garibaldi Provincial Park, B.C. (Osborn, et al., 2007)

4: Mt. Baker, North Cascades, W A (Osborn, et al., 2012)

4: Mt. Baker, North Cascades, WA (Kovanen & Slaymaker, 2005) No Data

5: Glacier Peak, North Cascades, WA (Beget, 1984, 1981) No Data

6: Enchantment Lakes Basin, North Cascades, WABilderbac ( k, 2004 ) 6: Enchantment Lakes Basin, North Cascades, W(AW aitt Jr, et al., 1982) No Data 7: Mt. Rainer, North Cascades, WA (Samolczyk, 2011)

7: Mt. Rainer, North Cascades, WA (Heine, 1998) No Data

8: Three Sisters Volcanoes, South Cascades, ORMa ( rcott et al., 2009) 9: Palisade Glacier, Sierra Nevada, CA (Bowerman & Clark, 2011) 10: Sierra Nevada, CA (Clark & Gillespie, 1997) No Data 11.7ka 9ka 6ka 3ka 1ka 0ka Key: Duration of Advance Maximum Age of Advance Minimum Age of Advance

FIGURE 3 – SUMMARIZED GLACIAL FLUCTUATION HISTORIES Ages of Holocene glacial advances from selected locations, corresponding to numbered blue stars on Figure 2. Note that numbers 1 and 2 represent the entire region, and thus do not have blue stars. N

SNOTEL site

Lower Basin

Upper Basin

Spider Gap

Chiwawa Mountain

Key:

FIGURE 4 – BASIN MORPHOLOGY OF THE LYMAN GLACIER Selected geomorphic features of the upper and lower basin of the Lyman Glacier. Note that Lyman Lake occupies the lower basin. Also note the cirque above the SNOTEL site. The horizontal distance from SNOTEL site to toe of glacier is ~3.5 km, and elevation is equal (1800 m). However, the respective aspects are di erent. Visible snow is typical of the annual minimum snow cover.

Imagery from Google Earth, capture date 6/23/2006, facing northwest at oblique angle. Spider Gap

Trim Line

Fluted Till

Terminal Moraine N

FIGURE 5 – GLACIAL FOREFIELD Selected geomorphic features of the glacial fore eld. Note that orange rocks de ning trim line have been carried downhill from source bedrock near Spider Gap.

Imagery from Google Earth, capture date 6/23/2006, facing southeast at oblique angle. Chiwawa Mountain

Spider Gap

FIGURE 6 – PHOTOGRAPH OF LYMAN GLACIER IN 1901 Photograph showing position of glacial terminus in relation to terminal moraine. Moraine appears as dark line behind proglacial lake in upper basin. There is no visible water or space between the moraine and the glacier. Note as well the pocket glacier in front of Chiwawa Mountain, which appears to have contributed meltwater to Lyman Lake.

Photograph taken by W.D. Lyman, 1901, facing south, from Cloudy Pass. Digital image taken from “The Columbia River: Its History, Its Myths, Its Scenery, Its Commerce” by W.D. Lyman, digitally created as a Project Gutenberg EBook (Lyman, 1909). N

Talus Slope Delta

Upper Lake Basin Moraine

Lower Lake Basin

Outlet

FIGURE 7 – LYMAN LAKE Selected geomorphic features of Lyman Lake. Note that the upper lake basin may receive sediment from the extensive talus slopes to the southwest. The outlet creek ows over bedrock.

Imagery from Google Earth, capture date 6/23/2006, facing southwest at oblique angle. 0 0.25 0.5 1 1.5 2 Kilometers Ü

FIGURE 8 – BEDROCK GEOLOGY OF THE LYMAN GLACIER BASIN Geologic map overlain on 10m DEM hillshade. Note the metamorphic rocks that compose Chiwawa Mountain and the Spider Gap area, in contrast to the plutonic rocks that surround the remainder of the basin.

Geologic map taken from Cater and Crowder's 1967 1:62,500 Holden quadrangle. FIGURE 9 - LYMAN LAKE SNOTEL SITE Photograph of Lyman Lake SNOTEL site, with snow pillow circled. Graph is a sample plot of Snow Water Equivalent (SWE) as measured by the snow pillow versus date. Note that the x-axis begins at October 1st, the start of the water year. The blue line represents the maximum historical SWE at the given date, the red line represents the minimum, and the green line is the historical average. The black line is the actual SWE recorded at the SNOTEL site in the current water year.

Photograph and graph taken from the National Oceanic and Atmospheric Administration’s Northwest River Forecast Center: Snow-Station Information webpage. Chiwawa Mountain

Spider Gap

Terminal Moraine

FIGURE 10 – PHOTOGRAPH OF LYMAN GLACIER IN 1915 Photograph showing glacier in early stages of 20th century retreat. The glacier’s volume occupies the entire valley, from wall to wall and nearly up to Spider Gap. Note the visible water behind the terminal moraine, indicating a substantial gap between the moraine and the glacial terminus. Also note the continuity of the glacier from Spider Gap to the cirque underneath Chiwawa Mountain. Arrow indicates ow from the gap. Finally, note the thickness of the glacier. The terminal moraine is ~ 5 m, for scale.

Photograph taken by L.D Lindsley, 1915, facing south. Digital image taken from University of Washington’s Special Collections. Chiwawa Mountain Spider Gap

Terminal Moraine

FIGURE 11 – PHOTOGRAPH OF LYMAN GLACIER IN 1921 Photograph showing source areas for Lyman Glacier. Besides the cirque directly above the main body of the glacier, signi cant snow is accumulating in snow elds beneath Spider Gap and Chiwawa Mountain. The modern descendants of these snow elds are visible at the base of Figure 4.

Photograph taken as part of The Mountaineers 1921 Glacier Peak Outing, facing south. Digital image taken from University of Washington’s Special Collections. Chiwawa Mountain Spider Gap

FIGURE 12 – PHOTOGRAPH OF LYMAN GLACIER IN 1997 Photograph showing the glacier occupying only its sheltered cirque after a century-long retreat. Note the snow elds that formerly owed into the glacier are now disconnected. The terminal moraine is hidden by the trees in the foreground.

Photograph taken by Ari Jumpponen, 1997, facing south. Digital image taken from Jumpponen et al. (1998). Spider Gap

Chiwawa Mountain A

B C FIGURE 13 (A-C) – RECESSION OF LYMAN GLACIER DURING THE 20TH CENTURY (A) shows the reconstructed outlines of the retreating Lyman Glacier from 1890 (blue line), 1921, 1958, 1979 and 2008. (B) shows the reconstructed horizontal retreat distance from the Little Ice Age terminal moraine. (C) shows the reconstructed decrease in volume (red) and areal extent (blue) over the 20th century retreat.

Modi ed gures taken from Pelto (2009). FIGURE 14 – CLIMATE CONDITIONS FOR MODERN GLACIERS Empirical plot showing temperature and precipitation conditions at 32 glaciers world-wide in the 1980s. The containing lines, or “envelope”, indicate favorable glacial climate conditions. If a modern glacier’s paleo-ELA is known, the modern climate conditions at the paleo-ELA can be plotted (red dot). The di erence between the modern climate at paleo-ELA and the envelope represent the amount of climate change necessary to make climate conditions at the paleo-ELA viable for a glacier to exist. Note that this does not provide a unique solution; rather, a range of temperature / precipitation conditions are possible, represented by the line connecting the blue dots.

Modi ed gure taken from Leonard (1989).