Properties of Glacial Ice and Glacier Classification

☆CglaFig

Tre

Luke Copland, Department of Geography, Environment and Geomatics, University of Ottawa, Ottawa, ON, Canada

© 2020 Elsevier Inc. All rights reserved.

1 Properties of glacial ice 12 Glacier classification schemes 23 Morphological classification 23.1 Unconstrained by topography 23.1.1 Ice sheet 23.1.2 Ice cap 23.1.3 Ice stream 33.1.4 Outlet glacier 43.1.5 Ice shelf 43.2 Constrained by topography 53.2.1 Icefield 53.2.2 Valley glacier 53.2.3 Cirque glacier 73.2.4 Piedmont glacier 73.2.5 Hanging glacier 74 Thermal classification 74.1 Temperate (“warm”) 84.2 Cold (“polar”) 94.3 Polythermal 95 Conclusions 10References 10

GlossaryAblation Any process which results in the loss of mass from a body of snow or ice. This typically occurs from surface melting,but can also occur due to processes such as wind redistribution, sublimation and iceberg calving.Accumulation Any process which results in the addition of mass to a body of ice or snow. This most commonly occurs fromsnowfall, but may also occur due to other processes such as avalanching from surrounding slopes, wind drifting or rainfalldirectly onto an ice mass.Cold ice (also referred to as polar ice) Ice which is at a temperature below the pressure melting point, and which thereforecontains little free liquid water.Firn Snow which has survived one summer melt season, but is not yet glacier ice. Typically found in the accumulation area ofglaciers.Ice age Any period in Earth’s history when there is permanent snow or ice on the Earth’s surface; we are therefore currently inan ice age. Glacial/interglacial periods refer to relatively colder/warmer periods within an ice age, with the last glacial maximumoccurring between approximately 26,500 and 19,000 years ago.Temperate ice (also referred to as warm ice) Ice which is at the pressure melting point, and therefore contains liquid water.

1 Properties of glacial ice

Glaciers (and their various forms, including ice sheets and ice caps) are composed of ice that is moving under its own weight, whichtypically requires the ice to be �30 m thick or more. Glacier ice is a crystalline solid and primarily originates from atmospherically-derived snow. At the highest elevations of a glacier, in the accumulation area, more snow falls annually than is lost by ablation, soover time the old snow gets buried and compacted by new snow above, causing it to increase in density. When snow has survivedmore than onemelt season it is termed firn. Firn contains interconnected pores that allow for the movement of air andmoisture, butas the firn continues to get buried and densifies, it finally reaches the pore close-off depth (Obbard et al., 2011). At this depth thepores become interconnected, which isolates air in individual bubbles, and defines the start of glacial ice. On glaciers with highannual snowfall and wet, dense snow packs (e.g., coastal Alaska), this metamorphism process occurs within just a few years, and the

hange History : September 2020. L Copland updated chapter title, Glossary with generally minor wording changes, added a new first section on “Properties ofcial ice,” text modified throughout with generally minor wording changes to improve clarity, updated references, updated figures 1, 2, 3, 6 and added newure 4

atise on Geomorphology https://doi.org/10.1016/B978-0-12-818234-5.00014-6 1

2 Properties of Glacial Ice and Glacier Classification

transition to glacial ice can happen at depths of a few tens of meters or less. In contrast, on glaciers in polar environments (e.g.,interior Antarctica), there is low annual snowfall and the dry, low density snowpack can take hundreds to thousands of years to turnto glacial ice, at depths of over a hundred meters (Cuffey and Paterson, 2010).

Newly fallen snow has a density of �50–200 kg m−3, with lower densities related to snow falling in colder and calmerconditions, and higher densities to snow falling in warmer and windier conditions. Density rapidly increases as new snow isburied, with firn typically having a density of 400–830 kg m−3 (Cuffey and Paterson, 2010). Pure glacier ice has a density of917 kg m−3, although this varies slightly based on factors such as bubble content, debris content, depth and ice temperature. Icecrystals have a hexagonal structure caused by the tetrahedral arrangement of H2O molecules that they are formed from. In glaciers,ice crystals typically grow in size with depth (i.e., pressure) as large grains consume small grains. The finest-grained ice layers containice crystals with amean grain size of�1 mm, although ranges of�2 mm to�2 cm are more common as depth increases (Gow et al.,1997). Ice crystals grow preferentially along their long axis (c-axis), and become organized in the direction of ice motion due to theprocesses of recrystallization and ice deformation.

Geomorphologically, ice masses take on a variety of forms that reflect their surrounding topography and climate, as well as theirinternal physical characteristics such as ice temperature, thickness and subglacial water flow. Permanent ice masses currently exist inall of the major mountain ranges in the world, as well as over extensive lowland regions close to the poles, covering a total of�10%of the Earth’s land surface (IPCC, 2019). This totals an area of�16,000,000 km2, with approximately 97% of this currently occupiedby the Greenland and Antarctic ice sheets (Benn and Evans, 2010).

2 Glacier classification schemes

To aid in the description and understanding of permanent ice masses, two main classification schemes have been developed. Thefirst of these is based on the surface morphology of an ice mass; in other words, the physical shape and geometry of an ice mass andhow it relates to the surrounding topography. The second classification scheme is based on the internal thermal characteristics of anice mass, which is controlled by factors such as the surrounding climate, geothermal heat production, and insulation properties ofthe overlying ice cover. The remainder of this chapter provides a review of these classification schemes, with a particular focus onproviding examples of ice masses that fall into different categories. This review covers permanent terrestrial ice masses, includingthose attached to land. For a multilingual technical definition of all ice types that occur in the ocean, atmosphere, on land, and inrelation to snow, the reader is referred to Armstrong et al. (1973). A good photo glossary of glacial features and landscapes is alsoavailable at the Swiss “Glaciers online” website (https://www.swisseduc.ch/glaciers, accessed June 16, 2020).

3 Morphological classification

The surface form of ice masses provides the classification scheme that has been traditionally used in glaciology, including byreference books ranging from the oldest (e.g., Agassiz, 1840) to more recent ones (e.g., Benn and Evans, 2010; Singh et al., 2011).Terrestrial ice masses fall into two primary categories within the morphological classification scheme: those unconstrained bysurrounding topography, and those constrained by it.

3.1 Unconstrained by topography

3.1.1 Ice sheetIce sheets are the largest class of ice mass, covering entire continents and subsuming almost all underlying topography (particularlyin their central regions). This includes entire mountain ranges such as the Gamburtsev Mountains in East Antarctica, which areburied by >3 km thick ice (Bell et al., 2011). Ice masses need to cover an area of >50,000 km2 to be classified as an ice sheet(Armstrong et al., 1973), and at present there are only two ice sheets on Earth: the Greenland and Antarctic ice sheets. Antarctica isthe thickest of these, with a maximum ice depth of �4.9 km (Fretwell et al., 2013), compared to a maximum depth of �3.3 km forthe Greenland Ice Sheet (Bamber et al., 2013). At the peak of the last glacial maximum, which lasted from approximately26,500–19,000 years ago (Clark et al., 2009), ice sheets covered almost 30% of the land surface of the planet. In the northernhemisphere this included the Laurentide Ice Sheet over most of eastern North America, the Cordilleran Ice Sheet over most ofwestern North America, and the Eurasian Ice Sheet over most of northern Europe.

3.1.2 Ice capIce caps are somewhat similar to ice sheets in that they cover almost all underlying topography, but are smaller in size(<50,000 km2). As ice caps still require a substantial ice thickness to cover topography, they most commonly occur in high latitudelocations such as Ellesmere Island (e.g., Agassiz Ice Cap), Devon Island (e.g., Devon Ice Cap; Fig. 1), Svalbard (e.g., Austfonna) andIceland (e.g., Vatnajökull). These ice caps typically reachmaximum ice thicknesses of�500–1000 m; for example, Dowdeswell et al.(2004) measured a maximum ice thickness of 880 m on Devon Ice Cap, and Björnsson and Pálsson (2008) reported a maximumthickness of �950 m for Vatnajökull.

Fig. 1 Radial flow pattern displayed by Devon Ice Cap in the Canadian High Arctic, based on speckle tracking of Radarsat-2 image pairs collected in winter 2012.Surface elevations and ice thicknesses are greatest near the slow-moving central ice dome, with velocities highest along topographically constrained outlet glaciers.White lines indicate ice divides. Modified from Van Wychen W, Burgess DO, Gray L, Copland L, Sharp M, Dowdeswell J, and Benham T (2014) Glacier velocities anddynamic ice discharge from the Queen Elizabeth Islands, Nunavut, Canada. Geophysical Research Letters 41(2): 484–490. https://doi.org/10.1002/2013GL058558.

Properties of Glacial Ice and Glacier Classification 3

A particularly unique feature of both ice caps and ice sheets is that they typically display a convex surface form, with highestelevations and thickest ice near the ice cap center, and progressive thinning toward the margins. These central high elevation regionshave been classified as “ice domes” (Benn and Evans, 2010). They typically display a radial flow pattern, with ice movementemanating in all directions from the ice dome, and flow rates increasing toward the ice cap or ice sheet margins (Fig. 1; Van Wychenet al., 2014). The low-velocity areas that form the boundary between ice flowing into different drainage basins are termed ice divides(white lines in Fig. 1).

3.1.3 Ice streamIce streams are defined as rapidly moving areas of an ice sheet that are bounded by slowly moving ice; they are thereforeunconstrained by surrounding surface topography. They most commonly occur in areas away from ice domes due to theirrequirement for rapid ice movement, with examples including the �700 km long Northeast Greenland Ice Stream (Joughin et al.,2001), and the Siple Coast ice streams that flow into the Ross Ice Shelf, Antarctica (Rignot et al., 2011). Velocities on ice streamsreach up to�1 km year−1, and they typically move at least an order of magnitude faster than the ice that surrounds them. This rapidchange of motion typically creates a narrow shear margin, which is characterized by intense surface crevassing. Field, modeling, andremote sensing evidence suggests that ice streams are typically warm-bedded (i.e., contain liquid water) and underlain by deformingsediment, whereas the slowly moving ice around them is cold-bedded and typically underlain by bedrock (Blankenshipet al., 2001).

There have been debates as to whether the location of ice streams is controlled by subglacial topography, but the evidence isambiguous. For example, Joughin et al. (2001) discovered that some areas of the Northeast Greenland Ice Stream have a subglacialtopographic expression, but in other locations under this ice stream there is little evidence of any subglacial channel. Bennett (2003)describes the Siple Coast ice streams as being the only contemporary examples of “pure” ice streams in Antarctica, as they lack anysubglacial topographic control. However, subglacial troughs have a major influence on ice stream flow elsewhere in Antarctica, suchas for the Ferrigno Ice Stream (Bingham et al., 2012). Adjacent ice streams have sped up and slowed down in the past in both theSiple Coast (Anandakrishnan and Alley, 1997; Bennett, 2003) and Weddell Sea (Vaughan et al., 2003) sectors of Antarctica, mostlikely driven by changes in subglacial water availability. For example, the Kamb Ice Stream (formerly Ice Stream “C”) appears tohave stagnated in the mid-nineteenth century (Anandakrishnan et al., 2001), whereas the Whillans Ice Stream (formerly Ice Stream“B”) appears to have started rapid flow �450 years ago after 400 years of stagnation (Hulbe and Fahnestock, 2004).

4 Properties of Glacial Ice and Glacier Classification

3.1.4 Outlet glacierOutlet glaciers are defined by Armstrong et al. (1973) as “a valley glacier which drains an inland ice sheet or ice cap and flowsthrough a gap in peripheral mountains”, and are somewhat similar to ice sheets and ice caps in that they drain ice from a central icedome. However, they are bounded by topographic valleys or troughs and therefore form a transitional ice mass between one that isunconstrained by topography versus one that is. On the large ice sheets, many outlet glaciers (e.g., Lambert Glacier, Antarctica;Kangerdlugssuaq Glacier, Greenland) originate as ice streams in their upper reaches. In combination, outlet glaciers and ice streamsaccount for the bulk of the discharge from ice sheets (e.g., up to 90% for Antarctica; Bamber et al., 2000), with outlet glaciersproviding the primary drainage for ice caps. As such, much recent research has focused on understanding the dynamics andvariability of outlet glaciers and ice streams since they appear to provide a major control on broad-scale ice sheet/ice cap stabilityand mass balance (e.g., Pritchard et al., 2009; Rignot et al., 2011, 2019; Moon et al., 2014).

3.1.5 Ice shelfIce shelves comprise permanent floating ice masses that are attached to the coast (Fig. 2), and typically form the seaward extensionof ice streams and outlet glaciers. This is particularly true in Antarctica where they comprise �75% of the coastline (Shepherd et al.,2018), with the majority of the discharge from the ice sheet passing through them. Antarctic ice shelves are some of the largestglaciological features on Earth, reaching >1000 km in length, 2500 m in depth, and covering a total area of >1,500,000 km2, some11% of the area of the Antarctic Ice Sheet (Fretwell et al., 2013; Schaffer et al., 2016). Ice shelves are also present along coastlines inthe northernmost reaches of the Arctic (Copland andMueller, 2017), but are typically orders of magnitude smaller than those in theAntarctic. For example, some small ice shelves appear to exist around the margins of Severnaya Zemlya and Franz Josef Land in theRussian High Arctic, although many have recently broken up (Dowdeswell, 2017). The largest concentration of Arctic ice shelvesoccurs along northern Ellesmere Island, Canada, where 13 ice shelves had a total area of 535 km2 in 2015 (Mueller et al., 2017).In some locations, ice shelves can build to a great enough thickness that they ground on underlying topography and create a dome-like structure that rises above the surrounding ice shelf. This feature is termed an ice rise, with an example being the Ward Hunt IceRise on northern Ellesmere Island (Braun, 2017).

In Antarctica, ice shelves are typically fed from upstream outlet glaciers and ice streams, although significant sources of massinput may also come from snow accumulation on their surface (particularly on very large ice shelves such as the Ross; Crary et al.,1962), as well as the basal freeze-on of ocean water (Bernales et al., 2017). In the Arctic, some ice shelves on northern EllesmereIsland receive input from surrounding glaciers (e.g., Milne Ice Shelf ), whereas others have built-up from the in situ thickening of seaice and surface snow accumulation with no known glacial input (e.g., Ward Hunt Ice Shelf ) (Jeffries, 2017). The majority of mass

Fig. 2 Milne Shelf, Ellesmere Island, Canada, July 2014. Note the almost flat surface topography. Source: Luke Copland.

Properties of Glacial Ice and Glacier Classification 5

loss from ice shelves occurs from the calving of ice at their termini; these large, freely-floating ice masses are termed icebergs in theAntarctic and ice islands in the Arctic, and are typically defined as having a height of >5 m above sea level (Armstrong et al., 1973;Canadian Ice Service, 2005). Icebergs also refer to the production of any floating ice from glaciers, whether in freshwater or theocean, and whether in the Arctic, Antarctic or elsewhere.

3.2 Constrained by topography

This category forms the second main group of ice masses classified by morphology, and comprises the bulk of ice masses that occurin mountainous regions around the world.

3.2.1 IcefieldAn icefield (also spelled ice field) forms the largest category of ice mass constrained by topography, and is somewhat similar to anice cap in that they can reach substantial thicknesses (e.g., >1000 m for most of the central Northern Patagonia Icefield; Gourletet al., 2016) and typically drain in a radial pattern from a high elevation region. They do not share the same dome-like surface as icecaps, however, as they have generally low relief surfaces dissected by areas of bedrock that reach above their surface as nunataks.Icefields typically form across broad high-elevation regions in areas such as Yukon/Alaska (e.g., St. Elias Icefield; Figs. 3 and 4),southern South America (Northern and Southern Patagonia Icefield) and Ellesmere Island (Prince of Wales Icefield, MansonIcefield).

3.2.2 Valley glacierA valley glacier consists of a moving ice mass which is constrained by surrounding topography and overlooked by rock cliffs(Sugden and John, 1976) (Figs. 4 and 5). Motion is required for an ice mass to be defined as a glacier, which typically means thatglaciers have a minimum thickness of �30 m. Thinner stationary ice will simply form a snow patch or ice patch. Typical motion onvalley glaciers varies from approximately 10–1000 m year−1 (although this can vary widely on both short and long timescales), withhighest values where the ice is relatively thick and steep, and in maritime areas where there is highest snow accumulation (Waechteret al., 2015). Surface slope provides one of the main drivers for ice motion, so to maintain movement valley glaciers typicallydecrease in surface elevation along their length. They commonly receive their source from icefields or cirques (Fig. 4), and areusually most clearly developed in the ablation area, where they can form substantial U-shaped valleys (Fig. 5). In places where

Fig. 3 St. Elias Icefield, Yukon, August 2016. Source: Luke Copland.

Fig. 4 Satellite image of the southeast portion of the St. Elias Icefield, Yukon, showing the connection between an icefield, valley glaciers, cirques and icefalls. Redbox indicates area of the inset image. Source: Sentinel 2, August 3, 2019.

Fig. 5 View up the temperate Grosser Aletschgletscher, Switzerland, August 2010. This ice mass forms the longest valley glacier in the European Alps at a totallength of �23 km. Note the hanging glaciers in the upper right corner of the image. Source: Luke Copland.

6 Properties of Glacial Ice and Glacier Classification

glaciers flow over bedrock obstacles they can become extremely steep and form icefalls, which are typically heavily crevassed on theirsurface.

The smallest glaciers are termed ice aprons and consist of thin snow and ice accumulations that adhere to the side of a mountainand can form the source of avalanches (Armstrong et al., 1973). Ice aprons typically occur at high altitudes near the tops of

Properties of Glacial Ice and Glacier Classification 7

mountains, and contrast with glacierets which occupy depressions on shallower terrain and commonly receive much of theiraccumulation from snow avalanching and wind drifting (Benn and Evans, 2010).

3.2.3 Cirque glacierCirque glaciers form in bowl-like depressions near the tops of mountains called cirques, which are typically characterized by a flatfloor and steep sides (Figs. 4 and 6). They can either be entirely contained within the cirque depression or provide mass todownstream features such as icefalls and valley glaciers, and are very common in areas of glaciated mountainous terrain. Theytypically have a preferred north to northeasterly orientation in the northern hemisphere (and vice versa in the southern hemisphere)due to shading from the sun and atmospheric circulation patterns (Evans, 1977).

3.2.4 Piedmont glacierPiedmont glaciers, also called piedmont lobes, form where constrained valley glaciers and outlet glaciers reach open lowland areasand are able to spread out laterally, whereby they take on a characteristic tear-drop or lobate shape due to the flow of the ice (Fig. 7).They typically occur at lower altitudes (i.e., in ablation areas), and usually form glacier termini. They are common in high latituderegions such as Ellesmere Island, but are relatively uncommon in steep mountainous areas such as the Andes, Alps and Rockies. TheMalaspina Glacier in SE Alaska forms the largest piedmont glacier in the world with a width of �65 km and area of �5000 km2

(including tributaries; Sauber et al., 2005)

3.2.5 Hanging glacierHanging glaciers occur on steep terrain inmountainous areas, and frequently form from ice that spills out of cirques or clings to rockcliffs (Fig. 5). They are typically small in area, and provide a frequent source of avalanches to regions below. In some cases, hangingglaciers can physically separate a glacier into an upper accumulation area and a lower ablation area fed almost entirely by iceavalanching from above.

4 Thermal classification

In addition to classifying ice masses based on their morphology, they can also be classified based on their thermal regime. Theinternal and basal temperature of ice masses can be determined directly from boreholes (e.g., Blatter, 1987), but borehole drilling istypically a time-consuming and expensive process. Many indirect measures are therefore used to determine internal and basal icetemperatures, including modeling (e.g., Gilbert et al., 2014), ground-penetrating radar (e.g., Björnsson et al., 1996), glacier

Fig. 6 Cirque glaciers and icefalls at an altitude of �3000 m in the St. Elias Mountains, Yukon. Source: Luke Copland.

Fig. 7 An example of polythermal piedmont glaciers issuing from small ice caps and outlet glaciers near the east coast of Ellesmere Island, Canada. Source: airphoto A16616-059 from the National Air Photo Library, Ottawa. © Her Majesty the Queen in Right of Canada.

8 Properties of Glacial Ice and Glacier Classification

hydrology and dynamics (e.g., Irvine-Fynn et al., 2011), and analysis of the glacial sedimentary record (e.g., Hambrey andGlasser, 2012).

The concept of pressure melting point is important when considering the thermal regime of glaciers, as the melting point ofwater decreases with increasing pressure. The depression in melting point occurs at a rate of 0.072 �C per million Pascals (Benn andEvans, 2010), which is equivalent to�110 m of ice. This means that the pressure melting point is�−2.5 �C at depths of�4000 m inthe deepest parts of the Antarctic Ice Sheet.

Knowledge of the thermal regime of ice masses is important as temperature strongly determines their dynamic behavior. Forexample, the viscosity of ice increases by a factor of five between −10 �C and −25 �C (Cuffey and Paterson, 2010). Ice masses whichhave a frozen bed typically undergo slow motion due to internal ice deformation alone. In contrast, ice masses at the pressuremelting point can have free water at their beds and typically experience higher surface velocities due to the addition of basalsediment deformation (where the ice lies on a soft bed) and basal sliding. Ice masses with beds at the pressure melting pointtypically exhibit high spatial and temporal variability in their surface motion in relation to variability in water inputs and associatedsubglacial water pressure (Willis, 1995). This means that variability in meltwater production may lead to changes in glacier thermalregime and associated ice motion (Phillips et al., 2010). It is clear that ice temperature is also important in controlling otherglaciological processes such as glacier surging (Fowler et al., 2001), subglacial sediment production (Hodson et al., 1997), andlandscape development (Dyke, 1993).

In general, ice masses are grouped into one of three different thermal classifications: temperate, cold or polythermal. Ice massesfrom the smallest glacierets to the largest ice sheets can exhibit any one (or more) of these temperature regimes.

4.1 Temperate (“warm”)

In this category, ice is at the pressure melting point throughout, except for a near-surface layer up to �20 m thick that is subject toseasonal temperature variations (Cuffey and Paterson, 2010). Water is able to freely exist within and at the bed of these ice masses,which means that they can undergo basal motion and move relatively rapidly. Temperate glaciers most commonly occur inrelatively low latitude and low altitude mountainous regions such as the Rockies, New Zealand, western Norway, and the EuropeanAlps (Fig. 5). Beyond the obvious requirement for a warm climate with an extended period of positive summer temperatures toremove any winter cold penetration, temperate glaciers also occur in regions with high mass turnover. This is because high snowfallrates tend to insulate underlying ice from cold winter temperatures, and high surface melt rates lead to latent heat release asmeltwater percolates into snow and refreezes. In addition, high ablation rates lead to the rapid removal of any winter build-up ofcold ice at low elevations. In the same climate regime, thicker glaciers tend to be warmer than thinner glaciers as ice acts as aninsulator that prevents cold surface temperatures from reaching the glacier bed.

Properties of Glacial Ice and Glacier Classification 9

4.2 Cold (“polar”)

In cold ice masses, the ice is below the pressure melting point throughout, meaning that water is not able to freely exist within themand at their base. This results in slow surface motion rates which are dominated by internal deformation, with little to no basalmotion. Cold glaciers are typically thin and occur in cold climates where there are low rates of mass turnover, such as the Dry Valleysof Antarctica. In the Dry Valleys mean annual temperatures are around −17 �C, the glaciers are frozen to their beds, and sublimationaccounts for up to 80% of their total annual ablation (Lewis et al., 1998)

4.3 Polythermal

Polythermal glaciers contain a mixture of both temperate and cold ice, and occur in high latitude and high altitude regionsthroughout the world. They can take on a variety of thermal structures, with free water only able to exist in their temperate parts(Blatter and Hutter, 1991). Fig. 8 sketches some of these thermal patterns, with types b and c occurring in relatively warmer climatescharacterized by sufficient melting at high altitudes to raise their accumulation areas to the pressure melting point due to latent heatrelease from the refreezing of percolating meltwater. Once warmed, this firn turns to ice and is carried to lower elevations by icemotion. It has a net vertical downward component in the accumulation area and an upward component in the ablation area. Thenear-surface of the ablation zone in these glaciers tends to be cold due to accumulation rates that are insufficient to insulate the icefrom winter cooling, and because water is not able to percolate the impermeable ice in these regions. Glacier types a and b aregenerally referred to as “predominantly warm polythermal,” and occur most commonly in polar maritime climates such as Svalbard(Björnsson et al., 1996) and inland areas of Scandinavia (e.g., Storglaciären, Sweden; Holmlund and Eriksson, 1989).

Glacier types d and e in Fig. 8 typically occur in colder, drier climates than types b and c. In these situations, there is insufficientpercolation and refreezing in the accumulation area to raise it to the melting point, so the cold ice from this region is carried to lowerelevations. In the ablation area the surface remains cold, but the bed is able to warm from geothermal heating, strain heating and theinput of surface meltwater via features such as crevasses and moulins. These types of glaciers are typically referred to as“predominantly cold polythermal,” and are common throughout regions such as the Canadian High Arctic (Blatter and Hutter,

Fig. 8 Various types of thermal structures within polythermal glaciers. Shading indicates temperate areas, white indicates cold areas, and dashed lines indicatemelting conditions along cold-temperate transitions. Arrows indicate approximate position of the equilibrium line altitude. Adapted from Blatter H and Hutter K (1991)Polythermal conditions in Arctic glaciers. Journal of Glaciology 37: 261–269, with permission of the International Glaciological Society.

10 Properties of Glacial Ice and Glacier Classification

1991). For example, John Evans Glacier on Ellesmere Island alternates between type d in the winter and type e in the summer, andcontains large volumes of subglacial meltwater across its terminus that can lead to rapid variations in horizontal and vertical motion(Copland et al., 2003). Glacier type a also falls into the category of predominantly cold polythermal, but would occur in warmerclimates than types d and e, where there is sufficient melting to raise the snow to the pressure melting point in the lower parts of theaccumulation area.

Recent studies of the Antarctic and Greenland ice sheets reveal that they have a more complex thermal structure than previouslybelieved, with cold surface temperatures but large areas of the bed at the pressure melting point. For example, modeling by Pattyn(2010) suggests that 55% of the grounded part of the Antarctic Ice Sheet is at the pressure melting point. They are thus best classifiedas being polythermal. Besides modeling, evidence for basal melting is provided by features such as the widespread occurrence ofsubglacial drainage (Wright et al., 2012; Moon et al., 2014), Antarctic subglacial lakes (Wright and Siegert, 2012), extensive basalfreeze-on in East Antarctica (Bell et al., 2011), and borehole drilling into temperate basal ice (Engelhardt et al., 1990; Bentley andKoci, 2007).

A few studies have demonstrated the existence of polythermal glaciers at high altitude in the Himalaya-Karakoram-Tibet region.For example, Miles et al. (2018) used borehole temperatures to show the presence of polythermal ice in the lower ablation area ofKhumbu Glacier, Nepal, with this part of the glacier consisting of �56% temperate ice and much of the remainder within 0.8 �C ofthe melting-point. Temperature measurements at depths of up to 92 m, on glaciers ranging in altitude from 3815 to 6303 m in Tibetand areas to the north, recorded mean temperatures of −6.6 to −11.3 �C (Liu et al., 2009). However, more measurements are neededto understand the distribution of internal and basal ice temperatures across this region, particularly because changes in thermalregime have been linked to changes in the occurrence of glacier surges in locations such as the Karakoram (Hewitt, 2005).

5 Conclusions

In the natural world, it should be remembered that ice masses exist in a wide variety of forms that do not always fall easily into thecategories described above. This broad continuum means that transitions occur between all of the categories listed here, and that asingle ice mass may fall into several different categories (particularly at different points along its length). For example, an ice particletraced from the center of Antarctica might pass through an ice sheet, ice stream, outlet glacier, and ice shelf before it reaches theocean. Even within a relatively short mountain glacier, an ice particle might pass from a cirque to an icefall to a valley glacier andthen to a piedmont glacier. Ice masses classified into one category by the morphological classification scheme can also be classifiedinto different categories based on the thermal scheme (and vice versa). For example, valley glaciers may be classified as temperate,cold or polythermal based on the climate regime that they exist in. Despite these complexities the classification of ice masses isimportant for aiding in their descriptions, understanding how they form, and providing knowledge about the controls on theirdynamics and changes.

References

Agassiz L (1840) Études sur les Glaciers, vol. 2. Neuchatel: H Nicolet.Anandakrishnan S and Alley RB (1997) Stagnation of ice stream C, West Antarctica by water piracy. Geophysical Research Letters 24: 265–268.Anandakrishnan S, Alley RB, Jacobel RW, and Conway H (2001) The flow regime of Ice Stream C and hypotheses concerning its recent stagnation. Antarctic Research Series

77: 283–294.Armstrong TE, Roberts B, and Swithinbank C (1973) Illustrated Glossary of Snow and Ice, 2nd edn. Cambridge: Scott Polar Research Institute. 60 pp.Bamber JL, Vaughan DG, and Joughin I (2000) Widespread complex flow in the interior of the Antarctic Ice Sheet. Science 287: 1248–1250. https://doi.org/10.1126/

science.287.5456.1248.Bamber JL, Griggs JA, Hurkmans RTWL, Dowdeswell JA, Gogineni SP, Howat I, Mouginot J, Paden J, Palmer S, Rignot E, and Steinhage D (2013) A new bed elevation dataset for

Greenland. The Cryosphere 7: 499–510.Bell RE, et al. (2011) Widespread persistent thickening of the East Antarctic Ice Sheet by freezing from the base. Science 331: 1592–1595. https://doi.org/10.1126/

science.1200109.Benn DI and Evans DJA (2010) Glaciers and Glaciation, 2nd edn. Oxford: Hodder Education.Bennett MR (2003) Ice streams as the arteries of an ice sheet: Their mechanics, stability and significance. Earth Science Reviews 61: 309–339.Bentley CR and Koci BR (2007) Drilling to the beds of the Greenland and Antarctic ice sheets: A review. Annals of Glaciology 47: 1–9.Bernales J, Rogozhina I, and Thomas M (2017) Melting and freezing under Antarctic ice shelves from a combination of ice-sheet modelling and observations. Journal of Glaciology

63(240): 731–744. https://doi.org/10.1017/jog.2017.42.Bingham RG, Ferraccioli F, King EC, Larter RD, Pritchard HD, Smith AM, and Vaughan DG (2012) Inland thinning of West Antarctic Ice Sheet steered along subglacial rifts. Nature

487: 468–471.Björnsson H and Pálsson F (2008) Icelandic glaciers. Jökull 58: 365–386.Björnsson H, Gjessing Y, Hamran S, Hagen JO, Liestol O, Pálsson F, and Erlingsson B (1996) The thermal regime of sub-polar glaciers mapped by multi-frequency radio-echo

sounding. Journal of Glaciology 42: 23–32.Blankenship DD, Morse DL, Finn CA, Bell RE, Peters ME, Kempf SE, Hodge SM, Studinger M, Behrendt JC, and Brozena JM (2001) Geologic controls on the initiation of rapid basal

motion for the West Antarctic ice streams: A geophysical perspective including new airborne radar sounding and laser altimetry results. Antarctic Research Series 77: 105–121.Blatter H (1987) On the thermal regime of an arctic valley glacier: A study of White Glacier, Axel Heiberg Island, N.W.T., Canada. Journal of Glaciology 33: 200–211.Blatter H and Hutter K (1991) Polythermal conditions in Arctic glaciers. Journal of Glaciology 37: 261–269.Braun C (2017) Chapter 6: The surface mass balance of the Ward Hunt Ice Shelf and Ward Hunt Ice Rise, Ellesmere Island, Nunavut, Canada. In: Copland L and Mueller DR (eds.) Arctic

Ice Shelves and Ice Islands, pp. 149–184. Dordrecht: Springer Nature.Canadian Ice Service (2005) MANICE: Manual of Standard Procedures for Observing and Reporting Ice Conditions, Rev. 9th edn. Ottawa: Environment Canada.

Properties of Glacial Ice and Glacier Classification 11

Clark PU, Dyke AS, Shakum JD, Carlson AE, Clark J, Wohlfarth B, Mitrovica JX, Hostetler SW, and McCabe AM (2009) The last glacial maximum. Science 325(5941): 710–714.https://doi.org/10.1126/science.1172873.

Copland L and Mueller D (eds.) (2017) Arctic Ice Shelves and Ice Islands. Dordrecht: Springer Nature. 425 pp. https://doi.org/10.1007/978-94-024-1101-0.Copland L, Sharp M, and Nienow P (2003) Links between short-term velocity variations and the subglacial hydrology of a predominantly cold polythermal glacier. Journal of Glaciology

49: 337–348.Crary AP, Robinson ES, Bennett HF, and Boyd WW (1962) Glaciological regime of the Ross Ice Shelf. Journal of Geophysical Research 67: 2791–2807.Cuffey KM and Paterson WSB (2010) The Physics of Glaciers, 4th edn. Oxford: Elsevier.Dowdeswell JA (2017) Ch. 3: Eurasian Arctic ice shelves and tidewater ice margins. In: Copland L and Mueller DR (eds.) Arctic Ice Shelves and Ice Islands, pp. 55–74. Dordrecht:

Springer Nature.Dowdeswell JA, Benham TJ, Gorman MR, Burgess D, and Sharp MJ (2004) Form and flow of the Devon Island Ice Cap, Canadian Arctic. Journal of Geophysical Research

109: F02002. https://doi.org/10.1029/2003JF000095.Dyke AS (1993) Landscapes of cold-centred Late Wisconsinan ice caps, Arctic Canada. Progress in Physical Geography 17(2): 223–247.Engelhardt H, Humphrey N, Kamb B, and Fahnestock M (1990) Physical conditions at the base of a fast moving Antarctic ice stream. Science 248: 57–59. https://doi.org/10.1126/

science.248.4951.57.Evans IS (1977) World-wide variations in the direction and concentration of cirque and glacier aspects. Geografiska Annaler: Series A, Physical Geography 59(3–4): 151–175.Fowler AC, Murray T, and Ng FSL (2001) Thermally controlled glacier surging. Journal of Glaciology 47(159): 527–538.Fretwell P, et al. (2013) Bedmap2: Improved ice bed, surface and thickness datasets for Antarctica. The Cryosphere 7: 375–393. https://doi.org/10.5194/tc-7-375-2013.Gilbert A, Gagliardini O, Vincent C, and Wagnon P (2014) A 3-D thermal regime model suitable for cold accumulation zones of polythermal mountain glaciers. Journal of Geophysical

Research Earth Surface 11999: 1876–1893.Gourlet P, Rignot E, Rivera A, and Casassa G (2016) Ice thickness of the northern half of the Patagonia Icefields of South America from high-resolution airborne gravity surveys.

Geophysical Research Letters 43(1): 241–249. https://doi.org/10.1002/2015GL066728.Gow AJ, Meese DA, Alley RB, Fitzpatrick JJ, Anandakrishnan S, Woods GA, and Elder BC (1997) Physical and structural properties of the Greenland Ice Sheet Project-2 ice core:

A review. Journal of Geophysical Research 102(C12): 26559–26575.Hambrey MJ and Glasser NF (2012) Discriminating glacier thermal and dynamic regimes in the sedimentary record. Sedimentary Geology 251–252: 1–33. https://doi.org/10.1016/j.

sedgeo.2012.01.008.Hewitt K (2005) The Karakoram anomaly? Glacier expansion and the “elevation effect,” Karakoram Himalaya. Mountain Research and Development 25: 332–340.Hodson AJ, Tranter M, Dowdeswell JA, and Gurnell AM (1997) Glacier thermal regime and suspended-sediment yield: A comparison of two high-Arctic glaciers. Annals of Glaciology

24: 32–37.Holmlund P and Eriksson M (1989) The cold surface layer on Storglaciären. Geografiska Annaler 71A: 241–244.Hulbe CL and Fahnestock M (2004) West Antarctic ice-stream discharge variability: Mechanism, controls and pattern of grounding-line retreat. Journal of Glaciology 50: 471–484.IPCC (2019) Summary for policymakers. In: Pörtner H-O, Roberts DC, Masson-Delmotte V, Zhai P, Tignor M, Poloczanska E,⋯Weyer N (eds.) IPCC Special Report on the Ocean and

Cryosphere in a Changing Climate. IPCC.Irvine-Fynn T, Hodson AJ, Vatne G, and Hubbard A (2011) Polythermal glacier hydrology: A review. Reviews of Geophysics 49(4), RG4002. https://doi.org/10.1029/2010RG000350.Jeffries MO (2017) Ch. 2: The Ellesmere ice shelves, Nunavut, Canada. In: Copland L and Mueller DR (eds.) Arctic Ice Shelves and Ice Islands, pp. 23–54. Dordrecht: Springer Nature.Joughin I, Fahnestock M, MacAyeal D, Bamber JL, and Gogineni P (2001) Observation and analysis of ice flow in the largest Greenland ice stream. Journal of Geophysical Research

106(D24): 34021–34024.Lewis KJ, Fountain AG, and Dana GL (1998) Surface energy balance and meltwater production for a Dry Valley glacier, Taylor Valley, Antarctica. Annals of Glaciology 27: 603–609.Liu Y, Hou S, Wang Y, and Song L (2009) Distribution of borehole temperature at four high-altitude alpine glaciers in Central Asia. Journal of Mountain Science 6: 221–227.Miles K, Hubbard B, Quincey D, Miles E, Sherpa T, Rowan A, and Doyle S (2018) Polythermal structure of a Himalayan debris-covered glacier revealed by borehole thermometry.

Scientific Reports 8: 16825.Moon T, Joughin I, Smith B, van den Broeke M, van de Berg WJ, Noël B, and Usher M (2014) Distinct patterns of seasonal Greenland glacier velocity. Geophysical Research Letters

41(20): 7209–7216.Mueller D, Copland L, and Jeffries MO (2017) Ch. 5: Changes in Canadian Arctic ice shelf extent since 1906. In: Copland L and Mueller D (eds.) Arctic Ice Shelves and Ice Islands,

pp. 109–140. Dordrecht: Springer Nature.Obbard RW, Baker I, and Lomonaco RW (2011) Firn. In: Singh VP, Singh P, and Haritashya UK (eds.) Encyclopedia of Snow, Ice and Glaciers. Dordrecht: Springer-Verlag.Pattyn F (2010) Antarctic subglacial conditions inferred from a hybrid ice sheet/ice stream model. Earth and Planetary Science Letters 295(3–4): 451–461. https://doi.org/10.1016/j.

epsl.2010.04.025.Phillips T, Rajaram H, and Steffen K (2010) Cryo-hydrologic warming: A potential mechanism for rapid thermal response of ice sheets. Geophysical Research Letters 37(20): L20503.

https://doi.org/10.1029/2010GL044397.Pritchard HD, Arthern RJ, Vaughan DJ, and Edwards LA (2009) Extensive dynamic thinning on the margins of the Greenland and Antarctic ice sheets. Nature 461: 971–975. https://

doi.org/10.1038/nature08471.Rignot E, Mouginot R, and Scheuchl B (2011) Ice flow of the Antarctic Ice Sheet. Science 333(6048): 1427–1430.Rignot E, Mouginot J, Scheuchl B, van den Broeke M, van Wessem MJ, and Morlighem M (2019) Four decades of Antarctic Ice Sheet mass balance from 1979–2017. Proceedings of

the National Academy of Sciences 116(4): 1095–1103.Sauber JM, Molnia B, Carabajal CC, Luthcke S, and Muskett R (2005) Ice elevations and surface change on the Malaspina Glacier, Alaska. Geophysical Research Letters 32: LL23S01.

https://doi.org/10.1029/2005GL023943.Schaffer J, Timmermann R, Arndt JE, Kristensen SS, Mayer C, Morlighem M, and Steinhage D (2016) A global, high-resolution data set of ice sheet topography, cavity geometry, and

ocean bathymetry. Earth System Science Data 8: 543–557. https://doi.org/10.5194/essd-8-543-2016.Shepherd A, Fricker HA, and Farrell SL (2018) Trends and connections across the Antarctic cryosphere. Nature 558: 223–232. https://doi.org/10.1038/s41586-018-0171-6.Singh VP, Singh P, and Haritashya UK (eds.) (2011) Encyclopedia of Snow, Ice and Glaciers. Dordrecht: Springer-Verlag.Sugden DE and John BS (1976) Glaciers and Landscape: A Geomorphological Approach. London: Edward Arnold.Van Wychen W, Burgess DO, Gray L, Copland L, Sharp M, Dowdeswell J, and Benham T (2014) Glacier velocities and dynamic ice discharge from the Queen Elizabeth Islands,

Nunavut, Canada. Geophysical Research Letters 41(2): 484–490. https://doi.org/10.1002/2013GL058558.Vaughan DG, Smith AM, Nath PC, and Le Meur E (2003) Acoustic impedance and basal shear stress beneath four Antarctic ice streams. Annals of Glaciology 36: 225–232.Waechter A, Copland L, and Herdes E (2015) Modern glacier velocities across the Icefield Ranges, St Elias Mountains, and variability at selected glaciers from 1959 to 2012. Journal of

Glaciology 61(228): 624–634. https://doi.org/10.3189/2015JoG14J147.Willis IC (1995) Intra-annual variations in glacier motion: A review. Progress in Physical Geography 19: 61–106.Wright A and Siegert M (2012) A fourth inventory of Antarctic subglacial lakes. Antarctic Science 24(6): 659–664.Wright AP, et al. (2012) Evidence of a hydrological connection between the ice divide and ice sheet margin in the Aurora Subglacial Basin, East Antarctica. Journal of Geophysical

Research Earth Surface 117: F01033. https://doi.org/10.1029/2011JF002066.