The Impact of Glacial Geomorphology on Critical Zone Processes

The Impact of Glacial Geomorphology on Critical Zone Processes

Chapter 12 The Impact of Glacial Geomorphology on Critical Zone Processes Kevin R. Gamache*, John R. Giardino*, Netra R. Regmi**, and John D. Vitek* ...

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Chapter 12

The Impact of Glacial Geomorphology on Critical Zone Processes Kevin R. Gamache*, John R. Giardino*, Netra R. Regmi**, and John D. Vitek* *

High Alpine and Arctic Research Program, Department of Geology and Geophysics, Texas A&M University, College Station, Texas, USA; **Department of Soil, Water and Environmental Sciences, University of Arizona, Tucson, Arizona, USA

12.1 INTRODUCTION The powerful erosional and depositional processes of glaciers have created much of the alpine, Arctic and Antarctic landscapes of Earth. Glaciers are one part of the five subsystems of Earth, namely the cryogenic regime or subsystem. Throughout geological time, glaciers have advanced and waned, but beginning in the Pleistocene, glaciers have had a massive effect on the present landscapes of many areas of North America, South America, Europe, Asia, and both Polar Regions. In addition to their sculpturing and depositing powers, glaciers have served as sinks for much of the freshwater on the planet, and in the past, covered as much as 30% of the land surface. In dealing with glacial environments, it is difficult to visualize this environment as part of the Critical Zone if we use the general definition of the Critical Zone, as extending from the top of the canopy to the bottom of the aquifer (NRC, 2001). Thus, it might be easy to dismiss or exclude the glacial environment as being part of the Critical Zone. In Polar Regions, no trees, shrubs or plants exist on glaciers, so where is the top of the canopy? We place the top of the canopy in the lower extent of the stratosphere. In alpine areas, numerous landforms make up the glacial environment. For example, various glacial deposits surround glaciers, often underlie glaciers, and even form atop glaciers. They have soil profiles and support plant life. Thus, we define the top of the canopy in this environment as extending from the top of trees, bushes, shrubs or grasses growing on various glacial deposits. What about the bottom of the Critical Zone? Although much of the terrain beneath glaciers is frozen, some areas below glaciers remain in an unfrozen state as the result of latitudinal location or pressure melting. Thus, water in the Developments in Earth Surface Processes, Vol. 19. http://dx.doi.org/10.1016/B978-0-444-63369-9.00012-6 Copyright © 2015 Elsevier B.V. All rights reserved.

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liquid state can and does exist. Lakes beneath Antarctic glaciers, for instance, also contain life. Complex interactions involving rock, soil, water, air, and living organisms exist. Glaciers themselves serve as stores of freshwater. The water supply for life sustaining resources in the City of Boulder, Colorado, for example, is the Arapaho glacier and its surrounding watershed. Nevertheless, it is water either in a frozen or liquid state. We have not found a definition that requires water to be in a liquid state to qualify as an aquifer. If we use the original definition of the Critical Zone as “… the heterogeneous, near surface environment in which complex interactions involving rock, soil, water, air and living organisms regulate the natural habitat and determine availability of life sustaining resources” (NRC, 2001, p. 2) then glacial environments are very much a part of the Critical Zone.

12.2  GOAL OF THIS CHAPTER Earth has a history of frequent local and global scale glaciations, which range from many hundreds of million years to recent. At the present time, most of the glaciers on Earth are located in Antarctica, ∼98% of which is covered by ice with an average thickness of 1.6 km; in the Arctic, the principal ice sheet that overlies Greenland has an area of 1,755,637 km2 and a thickness of up to 3,200 m; and individual valley glaciers, such as those of Vatnajökull, which is the largest glacier (∼1000-m thick) in Iceland (Gregory, 2010) (Fig. 12.1). Today, mountain glaciers, ice sheets and ice caps generate glacial processes that directly affect about 10% of the land surface of Earth, primarily in highlatitude regions (Gregory, 2010). This has not always been the case; however, in the past three million years more than 30% of the land surface was covered by glacial ice (Gregory, 2010). Earth was encapsulated in ice, around 750 M BP, a period referred to as Snowball Earth (Walker, 2003). Glaciers have ebbed and flowed since Snowball Earth ended. The glacial landscapes of today represent the most recently created landscapes on Earth. Although the Antarctic ice sheet has existed for at least 34 million years, extensive ice and permafrost did not exist prior to the Pliocene. At present 91.4% of the volume of glacier ice is in the Antarctic ice sheet and 8.3% is in Greenland, leaving just 0.3% for all the remaining glaciers in North America, South America, continental Europe, Asia and New Zealand (Gregory, 2010) (Fig. 12.2). Chris Rapley, former Director of the British Antarctic Survey, recognized these regions as being most sensitive to global change and providing a supersensitive early warning system for Earth (Rapley, 1999, 2006; Gregory, 2010). For that reason, understanding the basal and near-bed response of glaciers and ice sheets, and of glacier-bed interactions, is extremely important to glacial geomorphologists (Theakstone, 1982). Because glaciers do play a role in the Critical Zone, the major goal of this chapter is to provide detailed descriptions of processes and relationships in glacial areas, so that one can understand the important role these landscapes play in the Critical Zone. Although our focus

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FIGURE 12.1  Distribution of glaciers and ice around the world. Most of the world’s glacial ice is found in Antarctica and Greenland, but glaciers are found on nearly every continent. (Image/ photo courtesy of the National Snow and Ice Data Center, University of Colorado, Boulder.)

is on erosional and depositional processes of glaciers, glacier ecosystems, and hydrology, we link the processes with risks associated with glaciated regions and the response of glaciers to changes in climate.

12.3  GLACIER MASS BALANCE Today Critical Zone research is being driven by six overarching science questions. Of these six big-science questions, one that is directly related to glaciers, especially mass-balance studies, is focused on establishing in what manner theory and data can be linked from molecular to global scale to explain past transformations of the surface of Earth and predict the rate of development and planetary control of the Critical Zone (Banwart et al., 2012). Study and understanding of the processes and rates of growth and decline of glaciers is fundamental to answering the big science question. Glacier mass balance is expressed as the

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FIGURE 12.2  Approximate worldwide area covered by glaciers. These measurements were compiled by the World Glacier Monitoring Service, and published in the 2008 Publication, Global Glacier Changes: facts and figures (Roer et al., 2008). (Figure was adapted from National Snow and Ice Data Center, University of Colorado, Boulder.)

change in the volume of a glacier system by the gain or loss of ice on a temporal basis. The volume of a glacier is the function of the volume of ice it receives from snow accumulation and the amount of ice it loses by various processes, including sublimation and melting. Mass balance of a glacier can be described in terms of two processes as: accumulation and ablation of snow and ice. Snow and ice can be accumulated directly from precipitation, avalanches, windblown snow, and in minor amounts, from hoar frost, a layer of ice crystals formed by vapor transfer (i.e., sublimation followed by deposition) within dry snow beneath the snow surface. Accumulation can take place in the interior of the glacier when precipitation falls as rain and both meltwater and rainwater percolate through the snowpack and then refreeze. Accumulation can also take place at the base of a glacier or ice sheet by liquid-water freezing. The accumulated snow and ice can be transported downslope, as the glacier flows, until they reach a point where they are lost to the system, either by surface melting, surface meltwater runoff, sublimation, avalanching, wind erosion, and calving (the breakaway of iceblocks and icebergs). These processes are collectively known as ablation. Other processes of ablation include subaqueous melting, and melting within the ice and at the ice bed, which moves towards the terminus. Accumulation usually occurs over the entire glacier, but the rate of accumulation may change with elevation. Colder air temperatures at higher elevations of a glacier may result in more precipitation falling as snow, whereas, warmer air temperatures at lower elevations of a glacier may result in more precipitation falling as rain.

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The area where there is more accumulation than ablation (more mass gained than lost) is referred to as the accumulation zone. Similarly, ablation usually occurs over the entire glacier, but the part of the glacier that has more ablation than accumulation is the zone of ablation. Zones of accumulation and ablation are separated by a line where accumulation is equal to ablation and is referred to as the equilibrium line altitude (i.e., elevation). The snowline is commonly considered as the equilibrium line on many mountain glaciers. Glaciers losing more mass than they receive will be in a negative mass balance and so will be reduced in volume. The reduction of volume can result in the reduction of the thickness as well as the length of the glacier. Many mistakenly refer to this process as glacial retreat. The simple fact is glaciers do not retreat; they melt. Glaciers gaining more mass than they lose will be in positive mass balance and will grow and increase in volume. The glacier snout appears to moves backward in melting glaciers and the snout appears to move forward in expanding glaciers. Glaciers in equilibrium lose and gain approximately the same volume of snow and ice. They will neither increase nor decrease in depth and length. Accumulation and ablation of glacier ice vary over time and space, and depend on various factors including: elevation, aspect, relief, season, local and regional climate, latitude, and long-term change in climate (Kaser and Georges, 1999; Oerlemans, 2005; Oerlemans and Fortuin, 1992). Glaciers tend to be colder at the top than at the bottom. As a result, of this temperature profile, accumulation is greater in the upper extent of the glacier, and ablation is greater in the lower extent of the glacier. In the short-term, mass balance of a glacier varies throughout the year; glaciers typically receive more accumulation in the winter and experience more ablation in the summer (Dyurgerov and Meier, 1999). In the long-term, the dynamics of a glacier is primarily the function of change in climate (Haeberli and Beniston, 1998; Maisch, 2000). Study of glaciers around the world suggests that mass balances of glaciers have generally decreased during the past few decades, possibly as a result of current global warming (Oerlemans, 2005). Differences in the mass balance of a glacier due to changing climate causes adjustments in morphology to a new steady-state condition. The time it takes for the morphology of a glacier to adjust is known as its response time. The response time of a glacier depends on its mass-balance gradient, which is a critical factor controlled by the climatic regime in which the glacier is located. Mass-balance gradients measure the degree of change in net mass balance as a function of elevation. Net mass balance is generally at a maximum at the head of the glacier, zero at the equilibrium line, and minimum at the front or toe of the glacier. Glaciers with steeper mass-balance gradients have greater sensitivity to climate change than glaciers with less steep mass-balance gradients. Mass-balance gradients increase with increasing humidity; therefore, glaciers in wetter climates are more sensitive to changes in mass balance than glaciers in dry climates (Kuhn, 1981; Oerlemans and Fortuin, 1992). Glaciers in temperate regions, such as in the New Zealand Alps, receive very high amounts of

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precipitation, thus have a steeper mass-balance gradient (Chinn, 1996). These wet glaciers may have a shorter response time and higher climate sensitivity than cold and dry-polar glaciers that receive little accumulation and have relatively low rates of ablation (Rignot and Thomas, 2002). These glaciers may respond more slowly to climate change. The mass balance of a glacier over time can be measured by various techniques including: (1) installation of stakes, GPS and automatic weather stations, across a glacier (Andreassen et al., 2005; Van de Wal et al., 2008); (2) probing snowpits and crevasse stratigraphy (Pelto, 1990); and (3) analysis of time series remotely sensed images and aerial photographs (Bishop et al., 2004; Berthier et al., 2007). Stakes provide measurements of rates of accumulation and ablation at the stake locations on the glacier surface. This approach is time-­ consuming, often arduous, and thus challenging. Use of weather stations is key to understanding the fluxes of energy on the glacier, and the use of GPS is key to determining the movement dynamics of the glacier. A remote-sensing approach is generally the best approach for determining changes in the mass balance of a glacier because it is the most cost-effective approach and precludes commonly arduous fieldwork. However, the remote-sensing approach always requires ground validation of mass-balance measurements.

12.4  GLACIAL CHRONOLOGY AND QUATERNARY GLACIATION Earth has a history of frequent local and global scale glaciations at different time intervals and for various durations. Five major known ice ages have occurred in history (Hambrey and Harland, 1985). Ice ages are the periods of long-term reduction in the temperature of the surface of Earth and the atmosphere, which promotes expansion of continental and polar ice sheets and alpine glaciers. Within each ice age, a series of severe glacial and interglacial periods (more temperate) existed. The oldest unambiguous ice age in the history of Earth appears to be during the Proterozoic Eon in early Earth history (Table 12.1). Early Proterozoic (2400–2100 Ma). The Neoproterozoic glacial epoch (1000–544 Ma) represents the most severe ice age in the history of Earth (Hoffman et al., 1998; Rieu et al., 2007). During this time, Earth was fully or almost completely covered with ice, a period referred to as Snowball Earth (Walker, 2003). Earth is currently in an interglacial period of the Quaternary. The Quaternary Ice Age represents the last ∼2.5 Ma of geological time. Quaternary glaciation, more refined at the epoch stage as Pleistocene glaciation, is a series of glacial events separated by interglacial events. During this time, ice sheets expanded in Antarctica and Greenland and fluctuating ice sheets occurred in other parts of the world, also; a notable example being the Laurentide ice sheet (Huybrechts, 2002). Almost all of the present glacial landscapes worldwide are the result of Quaternary glaciation. Glacial activity during this

7th–8th

3rd–6th

2nd

1st

Nebraskan

Altonian

Günz– Mindel

Günz

Kansan

Yarmouthian

Mindel

Illinoian

Mindel– Riss

Sangamonian

Riss–Würm

Riss

Wisconsin

Würm

Beestonian

Cromerian

Anglian

Hoxnian

Wolstonian

Ipswichian

Devensian

Elburonian

Waalian

Elsterian

Holsteinian

Saalian

Eemian

Weichselian

Caracol

Rio Lico

Santa Maria

Valdivia

Llanquihue

Glacial period

Interglacial(s)

Glacial period(s)

Interglacial(s)

Glacial period

Interglacial

Glacial

Interglacial

MIS 12–478 ka

MIS 10–374 ka MIS 11–424 ka

MIS 6–191 ka

MIS 5e–123 ka

MIS 2–29 ka, near last glacial maximum MIS 3–57 ka MIS 4–71 ka MIS 5a–82 ka MIS 5b–87 ka MIS 5c–96 ka MIS 5d–109 ka

MIS 1–14 ka, end of the Younger Dryas marks the start of the Holocene, continuing to the present

Marine Isotope Stage (MIS)

621–676

MIS 15–621 ka MIS 16–676 ka

478–533–563 MIS 13–533 ka MIS 14–563 ka

424–478

374–424

130–200

115–130

12–71

12ka–9 ka

8 ka–5 ka

A summary of the global Quaternary glacial chronology and correlates each glacial/interglacial period with their respective Marine Isotope Stages. Derived from Ehlers and Gibbard (2008); Ehlers et al. (2011); and Frenzel (1992).

Pleistocene

Preboreal

Atlantic

5 ka–2.5 ka

Period 2.5 ka–1.50 a

Glacial/ interglacial period

Subboreal

South American

Subatlantic

Midwest US

Northern European

Holocene

Alps

British Isles

Anthropocene

Global region

Chronosequences/ glacial sequence

Epoch

TABLE 12.1 Global Quaternary Glacial Chronology

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period resulted in intense erosion and deposition of glacial sediments over large parts of the continents, modification of river systems, changes in sea level, creation of millions of glacial and pluvial lakes, climatic variability, and isostatic adjustment of the crust (Berger, 1978; Currey, 1990; Lambeck and Chappell, 2001). The last glacial period of the Quaternary ended approximately 11,000 years BP and currently Earth is in an interglacial period. Climatic warming caused the ice sheets from the last glacial period to begin to disappear. Remnants of the last glacial period, however, still exist in Greenland, Antarctica, and some alpine mountain regions. The current glaciation (i.e., Pleistocene glaciation) began at least ∼1.5 Ma in the Northern Hemisphere and continues to the Holocene (Jones, 2005). Evidence suggests climates in the Northern Hemisphere were glacial ∼70,000 BP, with two major short-term glacial periods ∼25,000 and ∼70,000 BP separated by a short period (<10,000 years) of climatic conditions associated with minimum glacial extent (Jones, 2005). Evidence further suggests that approximately one-third of the total land area of Earth was covered by ice at the Pleistocene glacial maximum. The largest ice sheet, known as the Laurentide of North America, covered the vast majority of Canada and the northern United States. The period from ∼130,000 BP to ∼70,000 BP is considered a period of interglaciation, and the period before that is considered glacial, which has slightly more ice coverage than the last glaciation (Jones, 2005). Minor glacial fluctuations have been reported for the past few hundred years. The period from 700BP to 100 BP has been termed the “Little Ice Age,” and was marked by observed glacier advances, especially in Europe (Hagen et al., 1993). Since then, however, most glaciers have shown a steady decline in extent.

12.5  GLACIAL FEATURES IN THE CRITICAL ZONE A glacier is a persistent mass of snow and ice that covers an area >0.1 km2, deforms under its own weight, shows evidence of downslope movement, and exhibits a fluid flow (Gregory, 2010). The mass of ice in a glacier forms by the recrystallization of snow or other forms of solid precipitation. Snow in the accumulation zone is transformed to firn (snow which has survived one summer melt season) and later to glacial ice where regrowth of ice crystals and elimination of air passages results in densities of 0.83–0.91 kg/m3. This may occur in only a year in valley glaciers, but can take several thousand years in ice-sheet environments (Gregory, 2010). Water and sediment are major components of glacial systems. The routes taken by the ice, commonly determine many of the resultant landscape features and landforms associated with the glacier environment. Sediment is carried on, in, and under the ice, and water can flow on, in, below, and around glaciers (Gregory, 2010). Sediment can be moved long distances from its source region.

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12.6  TYPES OF GLACIERS Glaciers exhibit a range of shapes, sizes, and morphology as a function of climate, topography, and location (Benn and Evans, 2014). For example glaciers range in size from a niche glacier to large ice sheets. Additionally, the movement mechanics and associated processes of mountain glaciers are mostly controlled by the topography. The movement mechanics and processes of ice sheets and ice caps are mostly independent of the topography. Therefore, for the first order classification, glaciers have been categorized based on the relation to the topography (Sugden and John, 1976). We here adopt this approach and classify glaciers as: (1) unconstrained by the topography and (2) constrained by the topography.

12.6.1  Glaciers Unconstrained by Topography 12.6.1.1  Ice Sheet and Ice Cap Ice sheets or ice caps, in the form of broad domes, are glaciers that bury the underlying topography with ice radiating outwards as a sheet. They have major patterns of ice flow that are largely independent to the bed topography (Gregory, 2010). The difference between an ice sheet and an ice cap is the area of coverage. Ice sheets generally cover >50,000 km2 and ice caps generally cover <50,000 km2 (Benn and Evans, 2014). Ice masses presently covering most of Antarctica and Greenland are larger than the threshold and are referred to as ice sheets, whereas ice masses in Nordaustlandet (Svalbard), Ellesmere Island, Baffin Island, and Iceland are mostly smaller than the threshold and are designated as ice caps (Benn and Evans, 2014). The Antarctic ice sheets cover around 98% of the continent (Laybourn-Parry et al., 2012). Ice sheets and ice caps form from the interplay of intensive accumulation of frozen water, mainly snow, and the insufficient melt during summer (Müller and Koch, 2012). Ice sheets of the northern hemisphere were characterized by the greatest dynamics known from about 120,000 BP down to about 10,000 BP. Several phases of ice-sheet growth and decay occurred during this time (Laybourn-Parry et al., 2012; Müller and Koch, 2012). These large masses of ice exert a tremendous weight on the underlying bedrock and cause it to be depressed. Upon melting, the weight gets removed and the land slowly rebounds, a process known as isostatic adjustment. The Hudson Bay area of Canada and the Baltic Sea between Finland and Norway are regions in which the rate of rebound has been measured. Although slow, the change signifies the impact that the presence of an ice mass had on the bedrock. Ice sheets and ice caps can be further categorized as ice domes, outlet glaciers, and ice streams. Ice domes are upstanding areas of ice sheets or ice caps where the movement of ice is relatively slow. Outlet glaciers and ice streams move rapidly out from the interiors of ice sheets and ice caps as channeled ice. These types of glaciers are constrained partially by the surrounding topography and exposed bedrock.

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12.6.1.2  Ice Shelves We do not discuss ice shelves in detail, as they are associated more with ocean areas adjacent to the Earth terrestrial environments that is, the Critical Zone.

12.6.2  Glaciers Constrained by Topography 12.6.2.1  Ice Field Ice fields develop in areas that are gentle in slope, but have locally fretted topography and sufficient elevation for the accumulation of ice. They appear similar to ice caps in terms of size, but differ in shape, and relation to the topography. They do not exhibit a dome-like structure, and the underlying surface topography influences the ice flow. Examples of well-known ice fields worldwide include the Columbia Icefields in the Canadian Rocky (Luckman et al., 1997), the ice fields of the St. Elias Mountains in the Canadian Yukon Territory/Alaska (Calkin, 1988), the Tien Shan/Kunlun Shan ice fields in China (Yi et al., 2002), and the Patagonian Ice Fields (Llibourty, 1998). Large valley glaciers drain all of these ice fields. 12.6.2.2  Valley Glaciers Valley glaciers form when ice is discharged from an ice field or a cirque into a deep bedrock valley. They mostly occur in mountain valleys in Polar Regions (Gregory, 2010) and alpine mountains, and may develop a simple, singlebranched planform, or a dendritic network (Penck, 1905). The form of the network and the glacial flow is commonly strongly influenced by the topography itself, the amounts of debris cover, and bedrock lithology and structure (Gregory, 2010). Bedrock slopes beneath valley glaciers are relatively steeper than the slopes beneath ice sheets and ice caps. In addition, the surrounding topography in valley glaciers is mostly ice-free over steepened slopes that generally feed snow and ice to the glacier surface (Benn and Evans, 2014) (Fig. 12.3). 12.6.2.3  Cirque Glaciers Cirque glaciers are partly enclosed by steep headwalls and may remain separated from the main valley glaciers or ice caps (Gregory, 2010). They form in bowl-shaped depressions, also known as bedrock hollows or cirques, located on the side of, or near mountains. They characteristically form by the accumulation of snow and ice avalanching from upslope areas. The size of cirque glaciers ranges from glaciers that are completely limited within hosting bedrock hollows, to glaciers that form the heads of large valley glaciers. The stability of a cirque glacier depends on various factors including, the size of the depression and the morphology of the surrounding topography, wind frequency and magnitude, and the availability of snow and ice in the cirque and surrounding slopes (Giardino et al., 1987; Sugden et al., 1999). The morphology of cirques and

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FIGURE 12.3  Blanca Peak (right center of photo) in the Sangre de Cristo Mountains, southcentral Colorado is 4372 m in elevation. Glaciated valleys and arêtes are prominently displayed in this scene. (Photo by John D. Vitek (October, 2005).)

bedrock hollows can shelter snow from wind, and if the snow persists through summer months, it becomes glacier ice. Snow may also be sheltered from wind on the leeward slope of a mountain and can be a major source of snow for the cirque glacier. Rock falls from upslope areas also play an important role in sheltering the snow and ice from sunlight. If enough rock falls onto the glacier, it may become a rock glacier (Giardino et al., 1987, 2014; Giardino and Vitek, 1988; Janke et al., 2013). If a cirque glacier advances far enough, it may become a valley glacier, and if the climate warms sufficiently to cause ablation, the valley glacier may form a cirque glacier.

12.6.2.4  Piedmont Glacier Piedmont glaciers develop by the process of valley glaciers debouching onto lowland areas after travelling through bedrock troughs and spreading out at the foot of mountain ranges. Some examples of piedmont glaciers include Malaspina Glacier in Alaska (Sharp, 1958) and Skeioararjokull glacier in Iceland (Sigurdsson, 1998). Many piedmont glaciers occur in the Canadian High Arctic, where subpolar glaciers debouch from plateau ice fields onto U-shaped valleys (Evans, 1990) (Fig. 12.4).

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FIGURE 12.4  Malaspina Glacier in southeastern Alaska is considered a classic example of a piedmont glacier. Piedmont glaciers occur where valley glaciers exit a mountain range onto broad lowlands, are no longer laterally confined, and spread to become wide lobes. (NASA Earth Observatory.)

12.6.2.5  Small-Sized Glaciers Masses of glacier ice also occur as ice fields or isolated glaciers. The smallest glacier ice-masses are known as ice aprons. They form by thin accumulations of snow and ice on mountainsides that is sometimes referred to as hanging glaciers. Similar patches of ice and snow occupying depressions on less precipitous terrain are commonly referred to as glacierets. They form as a result of snow drifting and avalanching from upslope areas. Glacierets that form because of ice avalanching from icefalls at steep plateau edges have been called fall glaciers. If a niche or rock bench in a mountain or valley side controls the location of an ice body, it is termed a niche glacier. Similar types of snow and ice accumulations in small depressions along coasts are referred to as ice fringes (Fig. 12.5).

12.7  EROSIONAL PROCESSES AND FORMS Although the body of literature concerning glacial-erosional processes and landforms is expanding, it is small compared to that concerning the processes and landforms of glacial deposition (Glasser and Bennett, 2004). The relationship of glacial erosional processes and landforms to former ice sheets determines the continuity of subglacial deformed layers, associated landforms, and sediments. Glacial erosion generates landscape evolution and relief production and moreover, can be used to assess paleoenvironmental factors (Glasser and Bennett, 2004).

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FIGURE 12.5  Arapaho Glacier, located along the Front Range, Colorado, in 2003. (Photo Courtesy of NASA Earth Observatory Program.)

12.8  EROSIONAL PROCESS Major processes of glacial erosion include: quarrying (also known as plucking), crushing abrasion, and mechanical and chemical erosion by glacial meltwater (Glasser and Bennett, 2004; Gregory, 2010). Glacial erosion creates a suite of landforms that are frequently observed in areas formerly occupied by ice sheets and glaciers (Glasser and Bennett, 2004). Quarrying involves the fracturing or crushing of bedrock beneath the glacier; and the entrainment of this fractured or crushed rock (Glasser and Bennett, 2004). Fracturing of bedrock occurs where a glacier flowing over bedrock creates pressure differences in the underlying rock, causing stress fields that are commonly sufficient to induce rock fracture (Glasser and Bennett, 2004; Morland and Morris, 1977). Plucking is particularly effective where a glacier flows over rock. Pressure exerted at the base of the ice melts the ice, which can refreeze in cracks in the underlying bedrock. Removing the pressure permits water to refreeze to the glacier and plucking occurs. Crushing occurs when the pressure exerted by basal rock fragments crushes the bedrock surface. Crushing develops crescentic fractures

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FIGURE 12.6  Exposed bedrock sculptured by glacial ice, Svartisen Glacier, Norway. (Photo by John D. Vitek (July 1988).)

called chattermarks, which indicate the direction of motion of the glacier. Abrasion involves the wearing down of rock surfaces by the grinding effect of rock fragments frozen into the base of glaciers. It occurs when bodies of subglacial sediment slide over bedrock (Glasser and Bennett, 2004), and produces smoothed bedrock surfaces that often exhibit parallel sets of scratches, called striations (1–10 mm diameter) and fine silt-sized particles (0.1 mm) known as rock flour (Fig. 12.6). The rate of glacial abrasion depends on ice velocity, presence of basal rock debris, hardness of the abrading rock fragment, removal of the rock flour, and thermal and pressure regime of the glacier (Boulton, 1976, 1982; Hallet, 1979, 1981; Schweizer and Iken, 1992). The rate of abrasion initially increases as glacial pressure and ice velocity increase, but decreases as the pressure becomes too great and ice at the base of the glacier begins to melt and release rock fragments from ice-rock fragment mixture. Abrasion increases with the increase in basal debris concentration. Furthermore, abrasion occurs only if the abrading rock fragments are harder than the bedrock. In addition, rock flour on the ice– rock bed interface needs to be flushed away by a constant supply of meltwater to sustain the process of abrasion. The thermal regime of a glacier also exerts a strong influence on the nature of erosion. The rate of abrasion beneath very cold glaciers (i.e., polar glaciers) tends to be negligible because of the lack of basal sliding; and in this condition, the adhesion of cold ice to bedrock, facilitates plucking (Benn and Evans, 2014). Once plucked, the blocks become the grinding tools that can cause abrasion.

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Meltwater erosion beneath ice sheets and glaciers may be caused by either mechanical or chemical processes (Glasser and Bennett, 2004). The effectiveness of meltwater as an erosional agent is dependent upon the susceptibility of the bedrock involved in erosion, the velocity and turbulent flow of the discharge regime, and the quantity of sediment in transport (Glasser and Bennett, 2004). The streams of meltwater that flow along the base of a glacier erode rock through the combined action of abrasion, hydraulic action, attrition, and solution (Hallet, 1981; Herman et al., 2011). Water at the base of a thick glacier generally remains under high hydrostatic pressure, which causes meltwater streams to have greater flow rates and erosive potential than the surface streams.

12.9  EROSIONAL FORMS Landforms developed by glacial erosion have a range of shapes, sizes, and morphology. We have categorized landforms developed by glacial erosion into three categories based on the sizes: (1) microscale forms; (2) intermediate scale forms; and (3) macroscale forms.

12.9.1  Micro-Scale Erosional Forms Major micro-scale forms of glacial erosion include striations, micro-crag and tails, bedrock gouges, and cracks. Striations are lines or scratches on rock surfaces, usually no more than a few millimeters in depth, produced by the process of glacial abrasion (Glasser and Bennett, 2004). These small grooves or scratches on bedrock surfaces, up to several meters long, are commonly associated with polished-bedrock surfaces (Glasser and Bennett, 2004). Striations demonstrate that ice sheets contained a significant amount of basal debris, experienced basal melting and a flow through basal sliding, and transported rock debris and sediments (Glasser and Bennett, 2004). Small tails of rock protected from glacial abrasion in the lee of resistant grains or mineral crystals on the surface of a rock are called micro-crag and tail. They are important for the reconstruction of former ice sheets because they provide clear evidence of the orientation and the direction of ice flow (Glasser and Bennett, 2004). Principal types of micro-scale cracks, gouges, and indentations created by glacial erosion on bedrock surfaces include: chattermarks, crescentic gouges, and crescentic cracks (Glasser and Bennett, 2004). Chattermarks and crescentic gouges generally consist of a shallow bedrock furrow with a crescentic outline (Glasser and Bennett, 2004). The convexity of the crescent is turned towards the direction of ice flow in crescentic gouges, whereas in chattermarks it is turned backwards (Glasser and Bennett, 2004). Chattermarks are commonly associated with larger bedrock grooves engraved on bedrock surfaces (Glasser and Bennett, 2004). Crescentic cracks, the vertical fracture of the rock without the removal of bedrock fragments, are normally curved in plane with the concavity turned towards the direction of flow (Glasser and Bennett, 2004).

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12.9.2  Intermediate Scale Landforms Intermediate scale landforms of glacial erosion, typically between 1 m and 1 km in size, have high relief on a wide range of scales, relatively smooth undulating surfaces, and upstanding streamline protrusions (Glasser and Bennett, 2004; Hindmarsh, 1999). These landforms may form parallel to the flow-direction of ice, and may be symmetric or asymmetric in down-glacier shape. These features have been referred to as roche moutonnées, rock drumlins, tadpole rocks and streamlined hills, ice-molded forms, whaleback and stoss and lee forms, and bedrock knolls depending on the size, shape, orientation with respect to the direction of flow, and morphology (i.e., asymmetry characteristics) (Fairchild, 1907; Flint, 1971, Dionne, 1987). Roche moutonnées are glaciated bedrock surfaces, usually in the form of rounded knobs, where the upstream and downstream slopes lie perpendicular to the general flow-direction of the former ice mass, and have a down-glacier asymmetry. The upstream side of a roche moutonnée forms a gentle, polished, and striated slope because of glacial scouring, and the downstream side resulted in a steep, irregular, and jagged slope because of glacial plucking. Drumlins are features similar to roche moutonnées, where an upstream side is steep and the downstream side is gentle, which is likely the result of the reduction in ice mass velocity (Benn and Evans, 2014). A crag and tail is an elongated, tapered ridge of till extending downstream formed commonly by the selective erosion of softer strata.

12.9.3  Large-Scale Erosional Forms Large-scale landforms developed by the glacial erosion include: glacial troughs, fjords, rock basins, knock and lochain, glacial lakes, and cirques. The glaciated landscape, particularly in mountains, exhibits distinguishing Ushaped (cross-section) valleys up to several thousand meters thick and tens of kilometers long; typically known as glacial troughs (Benn and Evans, 2014). They are the result of confining a glacier within valley walls, so that the glacier lacks uniform erosion and tends to deepen and widen the valley floor. This process transforms a commonly V-shaped stream valley into a more or less straight U-shaped valley because the U-shape provides relatively less frictional resistance to the relatively more viscous moving glacier (Boulton, 1982). The walls of U-shaped glacial valley may be almost vertical and striated by boulders dragged by the glacier. The valley floor may be covered with till or moraines. Because thickness of the ice is the dominant factor in the deepening process, smaller tributary glaciers erode the troughs less rapidly than the trough erosion of the main glacier. As a result the tributary troughs appear as valleys hanging on the wall of main glacial valley. When the glaciers melt, postglacial streams may form waterfalls from the mouths of the hanging valleys (Fig. 12.7).

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FIGURE 12.7  The U-shaped Huerfano River valley north of Blanca Peak, Sangre de Cristo Mountains, south central Colorado. (Photo by John D. Vitek (August 1978).)

Fjords, or fiords, are long narrow arms of the sea that commonly extend far inland. They result from marine inundation of glaciated valleys associated with rising sea levels. Fjords commonly are deeper in the middle and upper reaches than at the seaward end because of the greater erosional power of the glaciers closer to the source (Rafferty, 2011). Fjords are also U-shaped. The visible walls of fjords may rise vertically for hundreds of meters from the edge

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of the water, and close to the shore the water may be many hundreds of meter deep. The depth of a fjord may range up to more than a kilometer. For example, Sognefjord in Norway is ∼1308 m deep (Manzetti and Stenersen, 2010). The great depths of these submerged valleys are probably the result of the erosion from the thick glaciers that formed in these valleys.

12.10  GLACIAL TRANSPORT AND DEPOSITION Glaciers transport debris that has been plucked away or bits of rock that have been broken off, or fallen onto the glacier. Glacial debris can be transported in three ways: (1) on top of the glacier; (2) within the glacier; and (3) underneath the glacier. Glaciers deposit materials on the surfaces beneath, in surrounding areas, and at the margins, which become depositional features with a range of shapes, sizes and morphologies. Debris in the glacial environment may be deposited directly from the glacier ice (till) or from several associated processes involving glacial meltwater (outwash), and the resulting deposits are commonly termed as glacial drift. As the ice in a glacier moves from the zone of accumulation into the zone of ablation, it transports debris located beneath, within, and above the glacier toward its terminus or toward the outer margins where the ice velocity decreases. As the ice melts, the debris (till) that was originally frozen into the ice commonly forms a rocky and/or muddy blanket over the glacier margin. Typically, it is a nonstratified mixture of rock fragments and boulders in a fine-grained sandy or muddy matrix. Till occurs in various locations of a glacier, and depending on the mechanism of formation and location, can be categorized into different types. The deposit of glacial debris laid down more or less in place, as the ice melts, is called melt-out till or ablation till. In many cases, the glacial debris located between a moving glacier and its bedrock bed is severely sheared, compressed, and over-compacted; this type of deposit is called lodgement till. Tills often contain rock fragments and boulders that glaciers use to abrade the bedrock surface. If these rocks and boulders are different than the bedrock on which they are deposited, they are called erratics. Erratics are useful to determine the direction of ice movement and source region for the material (Fig. 12.8). Meltwater deposits, also called glacial outwash, can form in channels directly beneath or in front of the melting glacier or in lakes and streams in front of its margin. Outwash deposits generally consist of bedded, laminated, or stratified drift, with the individual layers composed of relatively well-sorted sediments (Benn and Evans, 2014). The grain size of individual deposits depends on the availability of different sizes of debris and also on the velocity of the current and the distance from the head of the stream. Larger boulders are deposited closer to the glacier margin, and the grain sizes of deposited material decreases with increasing distance from the glacier. The finest fractions, such as clay and silt, may be deposited in glacial lakes or ponds or transported all the way to the ocean.

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FIGURE 12.8  The massive sediment in the background is a poorly sorted till, deposited at the base of the now-thinning and retreating casement glacier. The boulders in the foreground are a lag deposit resulting from the removal of the finer grain-sizes in the till by the ice-marginal stream. Glacier Bay National Park, Alaska. USGS Photo.

12.10.1  Major Depositional Landforms of Valley Glaciers As a glacier moves downslope in a valley, it picks up rock debris from the valley walls and floor, and transports the sediment in, on, or under the ice. But, when the sediment–ice mixture reaches the lower parts of the glacier where ablation is the dominant process, the glacier deposits sediments along its margins during melting. The deposit is called a moraine. The size of the moraine depends on the size of the glacier, and the amount of material deposited by the glacier over time. Large valley glaciers are capable of forming moraines a few hundred meters high and many hundreds of meters wide (Bennett, 2001). If the position of the glacier margin is constant for an extended period of time, large moraines at the margin of the glaciers may develop because of the accumulations of larger amount of glacial debris (till). Depending on the location of formation with respect to the glacier, moraines can be categorized as end moraine, lateral moraine, and recessional moraine. Linear accumulations of till immediately in front or terminus of the glacier are called end moraines. The end moraine of the largest extent formed by the glacier during a period of given glaciation is referred to as the terminal moraine. Moraines formed along the valley slopes next to the side margins of the glacier are termed lateral moraines. The successive melting of a glacier can produce a series of recessional moraines; the joining of lateral moraines of two glaciers creates medial moraines (Figs 12.9 and 12.10).

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FIGURE 12.9  Complex moraine system adjacent to Fortuna Bay, South Georgia. (Photo by John D. Vitek (November 2013).)

FIGURE 12.10  Glacial tarn formation at the edge of a recessional moraine in a cirque near Grinnell Glacier, Glacier National 2013. (Photo by John D. Vitek.)

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Another major depositional landform of a valley glacier is a flute. Flutes are narrow, elongated, straight, parallel ridges of till which generally develop close to the lower margin of glaciers. They form when glaciers accumulate large volumes of debris beneath, as a result of the glacier gliding on a bed of pressurized muddy till. When basal ice flows around a bedrock knob or a boulder lodged in the substrate, a cavity can form in the ice on the lee side of the obstacle, and any pressurized mud present under the glacier may then be injected into this cavity and deposited as an elongate tail of till, or flute. The size of flutes range from a few centimeters to tens of meters in height and tens of centimeters to kilometers in length (Benn and Evans, 2014). The size of the flute mainly depends on the size of the obstacle and availability of subglacial debris. Flutes occur in valley and continental glaciers; however, very large flutes generally are also associated with continental ice sheets.

12.10.2  Depositional Landforms of Continental Glaciers Many depositional landforms associated with continental ice sheets are similar to the landforms of valley glaciers in terms of the formation. For example, the similar processes in both types of glaciers form terminal, end, and recessional moraines; however, the sizes associated with ice sheets are much larger. Morainic ridges of an ice sheet may extend for hundreds of kilometers, with hundreds of meters height and several kilometers width (Benn and Evans, 2014). In addition to moraine ridges, continental glaciers also deposit more or less continuous, thin (less than 10 m) sheets of till over large areas (Benn and Evans, 2014). Known as ground moraines, these deposits form undulating topography (hummocky) of low relief, with alternating small till mounds and depressions where swamps or lakes typically occur. Flutes are common features in areas covered by ground moraine. Continental glaciers develop streamlined, elongated mounds of sediment usually close to the edge of an ice sheet, called drumlins. Drumlins occur in groups of tens to hundreds, forming large drumlin fields (see Fig. 12.10, a map of the drumlins in central Wisconsin mapped by Atwood (1940). The long axis of a drumlin usually aligns parallel to the direction of regional ice flow. These features are generally asymmetric, where the stoss side is steeper than the lee side. Some drumlins consist entirely of till and display a fabric where long axes of the individual rocks and sand grains are aligned parallel to the direction of ice flow, whereas others have bedrock cores draped with till (Fig. 12.11).

12.10.3  Meltwater Deposits Much of the debris in the glacial environment of valley and continental glaciers is transported, reworked, and laid down by water. Whereas meltwater streams form glaciofluvial deposits, glaciolacustrine sediments accumulate at the margins and bottoms of glacial lakes and ponds.

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FIGURE 12.11  Map of drumlins showing the direction of ice motion in this portion of Wisconsin during the last advance of ice into this area (Atwood, 1940).

12.10.4  Glaciofluvial Deposits Meltwater streams at the snout of a valley glacier or along the margin of an ice sheet are generally laden with debris and have relatively high velocities. Beyond the glacier margin, once the walls of an ice tunnel no longer confine the water, it spreads out, and loses velocity. As a result, some of the load is deposited, which may cause the stream to separate into multiple channels separated by sand and gravel bar deposits. The process develops braided stream networks where the deposits usually consist of lenses of fine-grained, cross-bedded sands interbedded laterally and vertically with stringers of coarse, bouldery gravel referred to as outwash. In addition, the amount of sediment laid down by the stream generally (Fig. 12.12) decreases and becomes thinner with distance from the ice margin. The morphology of the outwash deposits depend on the surrounding topography. Where valleys are deep enough not to be buried by the glaciofluvial sediments, as is the case in most mountainous regions, the resulting elongate, planar outwash deposits are termed valley trains. In lowrelief areas, the deposits of several ice-marginal streams may merge to form outwash plains, or sandurs, which are wide plains. If the ice margin stabilizes at

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FIGURE 12.12  Braided channel flowing away from glacial snout creating a valley train ­deposit. Alaska, 2008. (Photo by John R. Giardino.)

a recessional position during glacial melting, additional valley trains or sandurs may form. The new sandur may lie at a lower elevation than the older sandur because of the downstream thinning of the outwash, and as a result, flat-topped remnants of the older plain may be left along the valley sides and develop terraces. Again, these structures can be used to reconstruct the positions of ice margins through time. Glaciofluvial streams that flow over the terminus of a glacier often deposit stratified drift in the channels and in depressions on the ice surface. These deposits, known as kames or kame moraines, commonly develop isolated mounds of bedded sands and gravels. Kames sometime form between the lateral margin of a glacier and the valley wall and develop terrace-like structure known as kame terraces. Streams also deposit stratified drift in subglacial or englacial tunnels, and as the ice melts away, these sinuous channel deposits may be left as long sinuous gravel ridges called eskers. The length can range from few hundred meters to hundreds of kilometers, and the height can range from few to tens of meters. Eskers have been used extensively as sources of sand and gravel for the construction industry. Steep-sided depressions, such as kettles, potholes, or ice pits are typical of many glacial and glaciofluvial deposits. Kettles form where till or outwash is deposited around ice blocks that are separated from active glacier because of ablation. When the ice melts, depressions are bordered by masses of glacial deposits. Lakes formed in kettles are called kettle lakes, and if a sandur or valley train contains kettles, it is referred to as a pitted outwash plain.

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12.10.5  Glaciolacustrine Deposits Glacial and proglacial lakes commonly occur on a glaciated landscape. Lakes on glaciated landscapes are formed by two major mechanisms: first, by the process of erosion, and second, by damming streams by ice or by glacial deposits, or by a combination of these processes. When a stream from a glacier enters a standing body of water, it deposits its bedload. The stream deposits coarser gravel and sand directly at the mouth of the stream as steeply inclined foreset beds. The finer, suspended silt and clay are transported farther into the lake, where they deposit as relatively flat-lying bottomset beds. As the sediment builds farther into the lake, the river deposits a thin veneer of subhorizontal gravelly topset-beds over the foreset units. The entire body of this deposit is comprised of foreset–topset complex. During warmer summer months, meltwater streams carry silt and clay into the lakes, and the silt settles out of suspension more rapidly than the clay. A thicker, silty summer layer is, thus, deposited. During winter, as the surface of the lake freezes and the meltwater discharge ceases, the clays contained in the lake water slowly settle out of suspension to form a thin winter clay layer. Such lacustrine deposits with annual silt and clay “couplets” are known as varves. These varves serve as good tools to date the age of deposit.

12.10.6  Glacier Hydrology Glaciers are dynamic systems. They store and mobilize snow, ice, water, and sediments, which characterize the overall system of glacial hydrology. Water on the surface of the glacier disappears through crevasses and holes in the glaciers, and flows emerge from the glacier snout. The hydrological system of a glacier can be categorized as: (1) supraglacial hydrology; (2) englacial hydrology; (3) subglacial hydrology; and (4) proglacial drainage systems. Supraglacial, or surface water on a glacier, is formed by the melting of ice (ablation) during the summer. The water flows off the glacier into crevasses and develops a network of channels, which are commonly sinuous where water can flow at rates of up to several meters per second (Cuffey and Paterson, 2010). Surface melt occurs in firn, the transitional state between snow and ice, and can pond above the impermeable ice. Saturation of the firn all the way to the surface creates a swamp zone where pools of standing water may form. The zone moves up glacier as the melt season progresses and the surface drains more rapidly as more ice is exposed, and the firn zone is filled with water (Cuffey and Paterson, 2010). Such processes of subglacial hydrology are considered responsible for large lakes on the surface of an ice sheet during the summer season in Greenland, and much of the coastal meltwater runoff in Antarctic coastal areas and ice shelves (Scambos et al., 2009). Englacial or within-glacier hydrological processes are controlled by crevasses, large tensional structures, and moulins, circular vertical shafts, which allow water to penetrate into the ice. Surface water cascades down into the

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ice sheets via crevasses and moulins and forms numerous water pockets and channels where water can remain trapped for some time (Fountain and Walder, 1998). Subglacial hydrological processes (hydrology underneath a glacier or ice sheet) are critically important in understanding the flow of Antarctic glaciers, where basal meltwater flows through the networks of large subglacial drainage, some of which extend up to hundreds of kilometers (Smith et al., 2007), and facilitate glacial erosion and movement of ice. Water in this system comes from two major sources: first by the basal melting of ice, and second by the downward percolation and flow of supraglacial and englacial water. Basal melting is primarily the result of geothermal heating and pressure of the ice mass. Water reaching the base of the ice sheets can be ponded in subglacial lakes, or flow through subglacial channels. Beneath the Antarctic ice sheet, these subglacial drainage channels are commonly connected to numerous subglacial lakes ­(Siegert et al., 2005; Smith et al., 2009). The proglacial area of temperate glaciers is characterized by abundant meltwater runoff from the glacier. The impoundment of proglacial meltwater in the deepened glacier basin may develop proglacial lakes. In addition, rapid exit of abundant meltwater from a glacier can form large braided-river plains, or sandurs, where glacial streams deposit, redeposit glacial sediments and rework glacial landforms. Examples of proglacial lakes occur in front of San Jose Glacier in James Ross Island, and braided streams are common characteristics in the northern Antarctic Peninsula (Carrivick et al., 2012) (Fig. 12.12).

12.11  GLACIER ECOSYSTEM Scientists have determined life exists in glaciers (Anesio and Laybourn-Parry, 2012; Kohshima, 1984). For example, a new species of cold-tolerant midge has been found in Himalayan glacier (Kohshima, 1984), microbial life has been known to occur in deep Greenland basal ice and Antarctica ice sheets (Price, 2000; Tung et al., 2006), bacteria and archaea have been found in the waters and sediments of Lake Whillans in Antarctic Ice sheet (Christner et al., 2014), and methanogenesis have been found in subglacial sediments from Robertson Glacier, Canadian Rockies (Boyd et al., 2010). Recent work suggests that moss plants can survive for centuries underneath glaciers (La Farge et al., 2013). Glacial surfaces receive sufficient sunlight, liquid water and direct contact with the atmosphere during summer. It has been suggested that temperatures far below zero do not present an absolute obstacle to microbial activity (Bakermans and Skidmore, 2011; Junge et al., 2006). These conditions create an environment favorable for the growth of life in three habitats including englacial habitat where rock-eating and other microbes live in thin films of water on the surface of debris entrained in the ice and the network of veins forming between individual ice crystals (Price, 2000; Tung et al., 2006), interface between ice

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and the atmosphere, and the surface of the ice. Keeping this in mind, snow that disappears at the peak of summer in the Northern Hemisphere (Nghiem et al., 2012), can contribute to the environment potential for a massive bloom of bacteria and microbes. In addition, summer conditions can be sufficiently gentle for some plants and animals to thrive on the ice surface. For example, mosses or ice worms (Porter et al., 2008) can increase the complexity of the icy food chain. Therefore, the, glacial ecosystem is one of the important topics for the study of the biological function in this Critical Zone. In addition, the topic is important because organic carbon trapped under the ice can be metabolized by microbes to form methane (Boyd et al., 2010), a potent greenhouse gas. Studies suggest that up to 21,000 petagrams (Pg) of organic carbon (10 times the permafrost carbon stock) might be trapped under the ice of Antarctica and that microbial conversion of this carbon to methane could be a major feedback in climate change (Wadham et al., 2012). In addition, subglacial microbes seem to greatly accelerate the weathering of minerals (Montross et al., 2013), and play a major role in the subglacial geochemical processes, and the glacier mass balance (Irvine-Fynn et al., 2012). In addition, ice algae (Yallop et al., 2012) and aggregates of microbes bound to minerals called cryoconite (Edwards et al., 2011) accelerate the rates of melting surface ice.

12.12  LIVING IN THE CRITICAL ZONE OF GLACIATED LANDSCAPE Glacial environments have significant relevance in global climate warming. Major components of these environments including snow, river and lake ice, sea ice, and frozen ground invoke positive feedback mechanisms that amplify global climate change and variability. For example, a decrease in snow and ice extent reduces the values of albedo and increases heat adsorption. Likewise, the thaw degradation of debris-covered glaciers may release large quantities of greenhouse gases (carbon and methane) to the atmosphere, which have an important feedback to climatic warming (e.g., Aleina et al., 2013; Schneider von Deimling et al., 2012; Vonk et al., 2013). Processes of erosion and deposition in glacial environments result in a unique landscape. Erosion by wind, nivation, frost weathering, meltwater (e.g., Thorn, 2009) will modify much of the Critical Zone environment. Seasonal patterns of stream discharge and sediment deposition in these environments are distinctive. Most of the high stream flows tend to occur during summer periods when snowmelt occurs (e.g., Mol et al., 2000; Vandenberghe, 2002). The shortlived peak discharges, in general, leave peculiar depositional features particularly in proglacial environments, such as poorly developed, shallow braided channels and large quantities of gravel and boulder deposits. Strong winds in glacial areas commonly move large quantities of loose sediment and soil (löess). The process dominantly occurs during (Fig. 12.13) dry summer months because

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FIGURE 12.13  Loess deposits form steep cliffs at the Troy L. Pewe Climatic Change Permafrost Reserve, Fairbanks, Alaska. (Photograph by J.R.Giardino (2008).)

sediments will dry out, meltwater from snow and ice decreases, which results in an increase in stream deposits available for aeolian transport. The process creates various landforms including sand dunes, löess accumulations, and sand sheets. Ventifacts, small rock fragments sculpted by sand blowing across a surface, are indicative of the role of wind in this environment (Knight, 2008). The Sand Hills of western Nebraska are the result of wind activity associated with Pleistocene glaciers. Natural processes in these environments can create high levels of risk for anthropogenic activities, including oil, gas, and mineral exploration, defense, tourism, transportation, and social and economic change. Most high-altitude and high-latitude cold environments on Earth remain covered by ice and snow in the winter, which then melts during the summer. Some of these events can be life-threatening and some affect the quality of life, infrastructure, natural resources, habitat, and agriculture. Most of the glacial landscapes in the world contain vast natural resources, notably hydrocarbons, gold, diamonds, iron, copper, and zinc as well as large reservoirs of water and construction materials. The exploitation of these resources in climatically rigorous and remotely located areas requires the knowledge of the processes, landforms, and potential hazards in the area to develop sound geotechnical, and engineering designs for constructions, including drilling, blasting, excavation of roads, buildings, bridges, pipelines, and other infrastructure.

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12.13  IMPLICATIONS FOR THE 21ST CENTURY In 2005, a mass of Antarctic ice, comparable to the size of California, briefly melted and separated from the main mass of ice, prompting the National Aeronautics and Space Administration (NASA) to report it as the most significant Antarctic melting in the past 30 years (Gregory, 2010). Such a phenomenon suggests that glaciers are one of the most dynamic systems on Earth. Glaciers are large, very dynamic, storage sites of water and sediment. They constantly exchange mass and energy with the atmosphere, hydrosphere, and other parts of the system of Earth. Glaciers gain mass by the accumulation of snow and ice on the surface, and lose mass by melting (ablation), iceberg calving, and many other processes. Therefore, the evolution of glacier mass depends on the balance between accumulation and ablation of snow and ice, which in turn depends on climate and local topographic factors (Kaser and Georges, 1999; Oerlemans, 2005; Oerlemans and Fortuin, 1992). Similarly glacier temperature evolves by the balance between energy input and output at the surface and the bed, and the heat generated within the glacier by ice flow. Both of these glacier characteristics strongly couple with the climate, and provide feedback responses to change in climate. Glacier mass balance and thermal structure impact many other parts of glacial and extra glacial environments at the local and global scale. For example, changes in snow and ice storage at the catchment-scale influence stream discharge and, thus, water resources and the risks from floods, whereas at the global scale, fluctuations in glacier volume directly impact the change in sea level (Bamber and Payne, 2004). Similarly, the fluctuation in glacier temperature, as a result of change in energy balance, influences many processes associated with glaciers, including surface ablation, englacial, and subglacial water flow, rates of glacier motion, and patterns of glacial erosion and deposition. For example, an increase in temperature increases the rates of snow and ice melt, subglacial flow, and glacial movements, which in turn increases the rates of erosion, deposition, and flooding. The processes of glaciers and ice sheets can also have significant influence on the atmosphere, in particular weather and climate modification at a local or global scale. Therefore, research on the links between glacier mass balance, thermal regime, and climate should be one of the prime focuses of research in the twenty-first century. Such research allows scientists to reconstruct long-term environmental changes from evidence of past glacier fluctuations, as well as future prediction of the environment in the glacial regions as a feedback response of current trends of global warming and change in climate. Fields of higher priority for research may include: (1) the better understanding of the influence of climate change in glacial and associated environments, to predict variety of potential natural events which may become hazardous, such as that associated with short- and long-term stream flow and water resources, seasonal and interannual climatic variability, local and global scale sea-level change; and (2) to make decisions on how to adapt to climate change.

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12.14 SUMMARY This chapter provides a quick overview of the processes that operate in glacial environments. These processes create a variety of landforms that modify local environments. Sediments are generated through erosion, moved to lower elevations and impact other geomorphic systems, especially the fluvial, aeolian and coastal systems. Whereas the amount of land impacted by glaciers continues to shrink (except in Antarctica and Greenland), the loss of ice means less water for human uses, including irrigation and industry. More water can also cause sea level to rise, which will impact every coastal city throughout the world. The economic impact of this rise will be staggering. To exclude and/or ignore the glacial environment as a component of the Critical Zone is a mistake that leads to humans being complacent in terms of learning how to respond to changes in the glacial system. As we continuously seek more resources to satisfy human needs, more exploration and development of glaciated regions will be undertaken. Without full knowledge of the processes and forms present in these areas, mistakes can be made that will ultimately increase the costs of goods and services. Throughout this chapter, we have linked the various processes with the potential inputs and how they can impact research in the Critical Zone such that human use of these regions can proceed efficiently.

REFERENCES Aleina, F.C., Brovkin, V., Muster, S., Boike, J., Kutzbach, L., Sachs, T., Zuyev, S., 2013. A stochastic model for the polygonal tundra based on Poisson-Voronoi diagrams. Earth Syst. Dynam. 4 (2), 187–198. Andreassen, L.M., Elvehøy, H., Kjøllmoen, B., Engeset, R.V., Haakensen, N., 2005. Glacier massbalance and length variation in Norway. Ann. Glaciol. 42 (1), 317–325. Anesio, A.M., Laybourn-Parry, J., 2012. Glaciers and ice-sheets as a biome. Trends Ecol. Evol. 27 (4), 219–225. Atwood, W.W., 1940. The Physiographic Provinces of North America. Ginn and Company, Boston. Bakermans, C., Skidmore, M., 2011. Microbial respiration in ice at subzero temperatures (−4 degrees C to −33 degrees C). Environ. Microbiol. Rep. 3 (6), 774–782. Bamber, J., Payne, T., 2004. Mass Balance of the Cryosphere: Observations and Modelling of Contemporary and Future Changes. Cambridge University Press, Cambridge; New York. Banwart, S., Menon, M., Bernasconi, S.M., Bloem, J., Blum, W.E., de Souza, D.M., Davidsdotir, B., Duffy, C., Lair, G.J., Kram, P., 2012. Soil processes and functions across an international network of Critical Zone Observatories: introduction to experimental methods and initial results. C.R. Geosci. 344 (11), 758–772. Benn, D., Evans, D.J., 2014. Glaciers and Glaciation. Routledge, London. Bennett, M.R., 2001. The morphology, structural evolution and significance of push moraines. Earth Sci. Rev. 53 (3), 197–236. Berthier, E., Arnaud, Y., Kumar, R., Ahmad, S., Wagnon, P., Chevallier, P., 2007. Remote sensing estimates of glacier mass balances in the Himachal Pradesh (Western Himalaya, India). Remote Sens. Environ. 108 (3), 327–338. Bishop, M.P., Olsenholler, J.A., Shroder, J.F., Barry, R.G., Raup, B.H., Bush, A.B.G., Copland, L., et al. 2004. Global Land Ice Measurements from Space (GLIMS): remote sensing and GIS investigations of the Earth’s cryosphere. Geocarto Int. 19 (2), 57–84.

392

Principles and Dynamics of the Critical Zone

Berger, A., 1978. Long-term variations of daily insolation and Quaternary climatic changes. J. Atmos. Sci. 35 (12), 2362–2367. Boulton, G., 1976. The origin of glacially fluted surfaces – observations and theory. J. Glaciol. 17, 287–309. Boulton, G.S., 1982. Processes and Patterns of Glacial Erosion. Springer, Netherlands. Boyd, E.S., Skidmore, M., Mitchell, A.C., Bakermans, C., Peters, J.W., 2010. Methanogenesis in subglacial sediments. Environ. Microbiol. Rep. 2 (5), 685–692. Calkin, P.E., 1988. Holocene glaciation of Alaska (and adjoining Yukon Territory, Canada). Quaternary Sci. Rev. 7 (2), 159–184. Carrivick, J.L., Davies, B.J., Glasser, N.F., Nyvlt, D., Hambrey, M.J., 2012. Late-Holocene changes in character and behaviour of land-terminating glaciers on James Ross Island. Antarctica J. Glaciol. 58 (212), 1176–1190. Chinn, T.J., 1996. New Zealand glacier responses to climate change of the past century. N. Z. J. Geol. Geophys. 39 (3), 415–428. Christner, B.C., Priscu, J.C., Achberger, A.M., Barbante, C.F., Carter, S.P., Christianson, K., Michaud, A.B., Mikucki, J.A., Mitchell, A.C., Skidmore, M.L., Vick-Majors, T.J., Team, W.S., 2014. A microbial ecosystem beneath the West Antarctic ice-sheet (vol. 512, p. 310, 2014). Nature 514 (7522), 394. Cuffey, K.M., Paterson, W.S.B., 2010. The Physics of Glaciers, fourth ed. Academic Press, Amsterdam, p. 704. Currey, D.R., 1990. Quaternary palaeolakes in the evolution of semidesert basins, with special emphasis on Lake Bonneville and the Great Basin, USA. Palaeogeogr. Palaeoclimatol. Palaeoecol. 76 (3), 189–214. Dionne, J.C., 1987. Tadpole rock (rocdrumlin): a glacial streamline moulded form. In: Rose, J., Menzies, J. (Eds.), Drumlin Symposium, Balkema, Rotterdam, pp. 149–159. Dyurgerov, M.B., Meier, M.F., 1999. Analysis of winter and summer glacier mass balances. Geogr. Ann. A Phys. Geogr. 81 (4), 541–554. Edwards, A., Anesio, A.M., Rassner, S.M., Sattler, B., Hubbard, B., Perkins, W.T., Young, M., Griffith, G.W., 2011. Possible interactions between bacterial diversity, microbial activity and supraglacial hydrology of cryoconite holes in Svalbard. ISME J. 5 (1), 150–160. Ehlers, J., Gibbard, P.L., Hughes, P.D., 2011. Quaternary Glaciations-Extent and Chronology: A Closer Look. Elsevier, Amsterdam. Ehlers, J., Gibbard, P.L., 2008. Quaternary glaciations extent and chronology. Evans, D.J., 1990. The last glaciation and relative sea level history of northwest Ellesmere Island, Canadian High Arctic. J. Quaternary Sci. 5 (1), 67–82. Fairchild, H.L., 1907. Drumlins of central New York. NY State Museum Bull. 111, 391–443. Flint, R.F., 1971. Glacial and Quaternary Geology. Wiley, New York, NY, p. 892. Fountain, A.G., Walder, J.S., 1998. Water flow through temperate glaciers. Rev. Geophys. 36 (3), 299–328. Frenzel, B., 1992. Atlas of Paleoclimates and Paleoenvironments of the Northern Hemisphere. Geographical Research Institute; Hungarian Academy of Sciences; Gustav Fischer Verlag, Budapest; Stuttgart; Jena; New York. Giardino, J., Regmi, N., Vitek, J., 2014. Rock Glaciers. In: Singh, V., Singh, P., Haritashya, U. (Eds.), Encyclopedia of Snow, Ice and Glaciers. Encyclopedia of Earth Sciences Series. Springer, Netherlands, pp. 943–948. Giardino, J.R., Shroder, J.F., Vitek, J.D. (Eds.), 1987. Rock Glaciers. Springer, London, p. 355. Giardino, J.R., Vitek, J.D., 1988. The significance of rock glaciers in the glacial-periglacial landscape continuum. J. Quaternary Sci. 3 (1), 97–103.

The Impact of Glacial Geomorphology on Critical Zone Processes Chapter | 12

393

Glasser, N.F., Bennett, M.R., 2004. Glacial erosional landforms: origins and significance for palaeoglaciology. Prog. Phys. Geogr. 28 (1), 43–75. Gregory, K.J., 2010. Earth’s Land Surface: Landforms and Processes in Geomorphology. SAGE Publications, London. Haeberli, W., Beniston, M., 1998. Climate change and its impacts on glaciers and permafrost in the Alps. Ambio 27, 258–265. Hagen, J.O., Liestøl, O., Roland, E., Jørgensen, T., 1993. Glacier atlas of Svalbard and Jan Mayen. Meddelelser NR. p. 129, Oslo. Hallet, B., 1979. A theoretical model of glacial abrasion. J. Glaciol. 23, 39–50. Hallet, B., 1981. Glacial abrasion and sliding: their dependence on the debris concentration in basal ice. Ann. Glaciol. 2 (1), 23–28. Hambrey, M.J., Harland, W.B., 1985. The Late Proterozoic Glacial Era. Palaeogeogr. Palaeocl. Palaeoecol. 51 (1–4), 255–272. Herman, F., Beaud, F., Champagnac, J.-D., Lemieux, J.-M., Sternai, P., 2011. Glacial hydrology and erosion patterns: a mechanism for carving glacial valleys. Earth Planet. Sci. Lett. 310 (3), 498–508. Hindmarsh, R.C.A., 1999. On the numerical computation of temperature in an ice-sheet. J. Glaciol. 45 (151), 568–574. Hoffman, P.F., Kaufman, A.J., Halverson, G.P., Schrag, D.P., 1998. A Neoproterozoic snowball earth. Science 281 (5381), 1342–1346. Huybrechts, P., 2002. Sea-level changes at the LGM from ice-dynamic reconstructions of the Greenland and Antarctic ice-sheets during the glacial cycles. Quaternary Sci. Rev. 21 (1), 203–231. Irvine-Fynn, T.D.L., Edwards, A., Newton, S., Langford, H., Rassner, S.M., Telling, J., Anesio, A.M., Hodson, A.J., 2012. Microbial cell budgets of an Arctic glacier surface quantified using flow cytometry. Environ. Microbiol. 14 (11), 2998–3012. Janke, J.R., Regmi, N.R., Giardino, J.R., Vitek, J.D. 2013, Rock Glaciers. In: Treatise on Geomorphology, Academic Press, San Diego, 238-273. Jones, P.N., 2005. Respect for the Ancestors: American Indian Cultural Affiliation in the American West. Bauu Institute, Colorado. Junge, K., Eicken, H., Swanson, B.D., Deming, J.W., 2006. Bacterial incorporation of leucine into protein down to-20 degrees C with evidence for potential activity in sub-eutectic saline ice formations. Cryobiology 52 (3), 417–429. Kaser, G., Georges, C., 1999. On the mass balance of low latitude glaciers with particular consideration of the Peruvian Cordillera Blanca. Geogr. Ann. A Phys. Geogr. 81 (4), 643–651. Knight, J., 2008. The environmental significance of ventifacts: a critical review. Earth Sci. Rev. 86 (1–4), 89–105. Kohshima, S., 1984. A novel cold-tolerant insect found in a Himalayan glacier. Nature 310 (5974), 225–227. Kuhn, M.C., 1981. Process and fundamental considerations of selected hydrometallurgical systems. Society of Mining Engineers of American Institute of Mining, Metallurgical, and Petroleum Engineers, New York, NY. La Farge, C., Williams, K.H., England, J.H., 2013. Regeneration of little ice age bryophytes emerging from a polar glacier with implications of totipotency in extreme environments. Proc. Natl. Acad. Sci. USA 110 (24), 9839–9844. Lambeck, K., Chappell, J., 2001. Sea level change through the last glacial cycle. Science 292 (5517), 679–686.

394

Principles and Dynamics of the Critical Zone

Laybourn-Parry, J., Tranter, M., Hodson, A.J., 2012. Ecology of Snow and Ice Environments. Oxford University Press, Oxford. Llibourty, L., 1998. Glaciers of Chile and Argentina. Geol. Surv. Prof. Pap. 1386, 1103. Luckman, B.H., Briffa, K.R., Jones, P., Schweingruber, F., 1997. Tree-ring based reconstruction of summer temperatures at the Columbia Icefield, Alberta, Canada, AD 1073–1983. Holocene 7 (4), 375–389. Maisch, M., 2000. The long-term signal of climate change in the Swiss Alps: Glacier retreat since the end of the Little Ice Age and future ice decay scenarios. Geogr. Fis. Dinam. Quat. 23, 139–151. Manzetti, S., Stenersen, J.H.V., 2010. A critical view of the environmental condition of the Sognefjord. Mar. Pollut. Bull. 60 (12), 2167–2174. Mol, J., Vandenberghe, J., Kasse, C., 2000. River response to variations of periglacial climate in mid-latitude Europe. Geomorphology 33 (3–4), 131–148. Montross, S.N., Skidmore, M., Tranter, M., Kivimaki, A.L., Parkes, R.J., 2013. A microbial driver of chemical weathering in glaciated systems. Geology 41 (2), 215–218. Morland, L., Morris, E., 1977. Stress in an elastic bedrock hump due to glacier flow. J. Glaciol. 18, 67–75. Müller, J.A., Koch, L., 2012. Ice-sheets: Dynamics, Formation and Environmental Concerns. Earth Sciences in the 21st Century. Nova Science Publisher’s, Inc., Hauppauge, NY, 212 p. National Research Council, 2001. Committee on Basic Research Opportunities in the Earth. Basic Research Opportunities in Earth Science. 0-309-07133-X, National Research Council, Washington, DC. Nghiem, S.V., Hall, D.K., Mote, T.L., Tedesco, M., Albert, M.R., Keegan, K., Shuman, C.A., DiGirolamo, N.E., Neumann, G., 2012. The extreme melt across the Greenland ice-sheet in 2012. Geophys. Res. Lett. 39, 39. Oerlemans, J., 2005. Extracting a climate signal from 169 glacier records. Science 308 (5722), 675–677. Oerlemans, J., Fortuin, J.P., 1992. Sensitivity of glaciers and small ice caps to greenhouse warming. Science 258 (5079), 115–117. Pelto, M. (1990). Annual balance of North Cascade, Washington glaciers predicted from climatic records. Eastern c, 201. Penck, A., 1905. Glacial features in the surface of the Alps. J. Geol. 13 (1), 1–19. Porter, P.R., Evans, A.J., Hodson, A.J., Lowe, A.T., Crabtree, M.D., 2008. Sediment-moss interactions on a temperate glacier: Falljokull. Iceland. Ann. Glaciol. 48, 25–31. Price, P.B., 2000. A habitat for psychrophiles in deep Antarctic ice. Proc. Natl Acad. Sci. USA 97 (3), 1247–1251. Rafferty, J.P., 2011. Landforms. Britannica Educational Publishing, Chicago. Rapley, C., 1999. Invited keynote address: global change and the polar regions. Polar Res. 18 (2), 117–118. Rapley, C., 2006. The Antarctic Ice Sheet and Sea Level Rise. Avoiding Dangerous Climate Change. Cambridge University Press, Cambridge, pp. 25–27. Rieu, R., Allen, P.A., Plotze, M., Pettke, T., 2007. Climatic cycles during a Neoproterozoic “snowball” glacial epoch. Geology 35 (4), 299–302. Rignot, E., Thomas, R.H., 2002. Mass balance of polar ice sheets. Science 297 (5586), 1502–1506. Roer, I., Zemp, M., van Woerden, J., 2008. Global glacier changes: facts and figures. UNEP/Earthprint. Scambos, T., Fricker, H.A., Liu, C.C., Bohlander, J., Fastook, J., Sargent, A., Massom, R., Wu, A.M., 2009. Ice shelf disintegration by plate bending and hydro-fracture: satellite ­observations

The Impact of Glacial Geomorphology on Critical Zone Processes Chapter | 12

395

and model results of the 2008 Wilkins ice shelf break-ups. Earth Planet. Sci. Lett. 280 (1–4), 51–60. Schneider von Deimling, T., Meinshausen, M., Levermann, A., Huber, V., Frieler, K., Lawrence, D.M., Brovkin, V., 2012. Estimating the near-surface permafrost-carbon feedback on global warming. Biogeosciences 9 (2), 649–665. Schweizer, J., Iken, A., 1992. The role of bed separation and friction in sliding over an undeformable bed. J. Glaciol. 38, 77–92. Sharp, R.P., 1958. Malaspina Glacier, Alaska. Geol. Soc. Am. Bull. 69 (6), 617–646. Siegert, M.J., Carter, S., Tabacco, I., Popov, S., Blankenship, D.D., 2005. A revised inventory of Antarctic subglacial lakes. Antarctic Sci. 17 (3), 453–460. Sigurdsson, O., 1998. Glacier variations in Iceland 1930–1995. Jokull 45, 3–26. Smith, A.M., Murray, T., Nicholls, K.W., Makinson, K., Aoalgeirsdottir, G., Behar, A.E., Vaughan, D.G., 2007. Rapid erosion, drumlin formation, and changing hydrology beneath an Antarctic ice stream. Geology 35 (2), 127–130. Smith, B.E., Fricker, H.A., Joughin, I.R., Tulaczyk, S., 2009. An inventory of active subglacial lakes in Antarctica detected by ICESat (2003–2008). J. Glaciol. 55 (192), 573–595. Sugden, D.E., John, B.S., 1976. Glaciers and Landscape: A Geomorphological Approach. Edward Arnold, London. Sugden, D.E., Summerfield, M.A., Denton, G.H., Wilch, T.I., McIntosh, W.C., Marchant, D.R., Rutford, R.H., 1999. Landscape development in the Royal Society Range, southern Victoria Land, Antarctica: stability since the mid-Miocene. Geomorphology 28 (3), 181–200. Theakstone, W.H., 1982. Glacial Geomorphology. Prog. Phys. Geogr. 6 (2), 261–274. Thorn, E.C., 2009. Holocene microweathering rates and processes on ice-eroded bedrock, Roldal area, Hardangervidda, southern Norway. Geological Society, London, Special Publications January 1, 2009, vol. 320, pp. 29–49. Tung, H.C., Price, P.B., Bramall, N.E., Vrdoljak, G., 2006. Microorganisms metabolizing on clay grains in 3-km-deep Greenland basal ice. Astrobiology 6 (1), 69–86. Van de Wal, R.S.W., Boot, W., Van den Broeke, M.R., Smeets, C.J.P.P., Reijmer, C.H., Donker, J.J.A., Oerlemans, J., 2008. Large and rapid melt-induced velocity changes in the ablation zone of the Greenland ice sheet. Science 321 (5885), 111–113. Vandenberghe, J., 2002. Periglacial sediments: do they exist? Geological Society, London, Special Publications January 1, 2011, vol. 354, pp. 205–212. Vonk, J.E., Mann, P.J., Dowdy, K.L., Davydova, A., Davydov, S.P., Zimov, N., Spencer, R.G.M., Bulygina, E.B., Eglinton, T.I., Holmes, R.M., 2013. Dissolved organic carbon loss from Yedoma permafrost amplified by ice wedge thaw. Environ. Res. Lett. 8 (3), 9, 035023. Wadham, J.L., Arndt, S., Tulaczyk, S., Stibal, M., Tranter, M., Telling, J., Lis, G.P., Lawson, E., Ridgwell, A., Dubnick, A., Sharp, M.J., Anesio, A.M., Butler, C.E.H., 2012. Potential methane reservoirs beneath Antarctica. Nature 488 (7413), 633–637. Walker, G., 2003. Snowball Earth. Three Rivers Press, New York, p. 269. Yallop, M.L., Anesio, A.M., Perkins, R.G., Cook, J., Telling, J., Fagan, D., MacFarlane, J., Stibal, M., Barker, G., Bellas, C., Hodson, A., Tranter, M., Wadham, J., Roberts, N.W., 2012. Photophysiology and albedo-changing potential of the ice algal community on the surface of the Greenland ice-sheet. ISME J. 6 (12), 2302–2313. Yi, C., Li, X., Qu, J., 2002. Quaternary glaciation of Puruogangri – the largest modern ice field in Tibet. Quaternary Int. 97, 111–121.