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21.2: Glaciers - Geosciences

21.2: Glaciers - Geosciences


A glacier (US /ˈɡleɪʃər/ or UK /ˈɡlæsiə/) is a persistent body of dense ice that is constantly moving under its own weight; it forms where the accumulation of snow exceeds its ablation (melting and sublimation) over many years, often centuries. Glaciers form only on land and are distinct from the much thinner sea ice and lake ice that form on the surface of bodies of water.

The Baltoro Glacier in the Karakoram, Baltistan, Northern Pakistan. At 62 kilometres (39 mi) in length, it is one of the longest alpine glaciers on earth.

Ice calving from the terminus of the Perito Moreno Glacier in western Patagonia, Argentina

The Aletsch Glacier, the largest glacier of the Alps, in Switzerland

The Quelccaya Ice Cap is the largest glaciated area in the tropics, in Peru

On Earth, 99% of glacial ice is contained within vast ice sheets in the polar regions, but glaciers may be found in mountain ranges on every continent except Australia, and on a few high-latitude oceanic islands. Between 35°N and 35°S, glaciers occur only in the Himalayas, Andes, Rocky Mountains, a few high mountains in East Africa, Mexico, New Guinea and on Zard Kuh in Iran.[1]

Glacial ice is the largest reservoir of freshwater on Earth.[2] Many glaciers from temperate, alpine and seasonal polar climates store water as ice during the colder seasons and release it later in the form of meltwater as warmer summer temperatures cause the glacier to melt, creating a water source that is especially important for plants, animals and human uses when other sources may be scant. Within high altitude and Antarctic environments, the seasonal temperature difference is often not sufficient to release meltwater.

Because glacial mass is affected by long-term climate changes, e.g., precipitation, mean temperature, and cloud cover, glacial mass changes are considered among the most sensitive indicators of climate change and are a major source of variations in sea level.

Etymology and related terms

The word glacier comes from French. It is derived from the Vulgar Latin glacia and ultimately from Latin glacies meaning “ice”.[3] The processes and features caused by glaciers and related to them are referred to as glacial. The process of glacier establishment, growth and flow is called glaciation. The corresponding area of study is called glaciology. Glaciers are important components of the global cryosphere.

Types

Main article: Glacier morphology

Mouth of the Schlatenkees Glacier near Innergschlöß, Austria

Glaciers are categorized by their morphology, thermal characteristics, and behavior. Alpine glaciers, also known as mountain glaciersor cirque glaciers, form on the crests and slopes of mountains. An alpine glacier that fills a valley is sometimes called a valley glacier. A large body of glacial ice astride a mountain, mountain range, or volcano is termed an ice cap or ice field.[4] Ice caps have an area less than 50,000 km² (20,000 mile²) by definition.

Glacial bodies larger than 50,000 km² are called ice sheets or continental glaciers.[5] Several kilometers deep, they obscure the underlying topography. Only nunataks protrude from their surfaces. The only extant ice sheets are the two that cover most of Antarctica and Greenland. They contain vast quantities of fresh water, enough that if both melted, global sea levels would rise by over 70 meters.[6]Portions of an ice sheet or cap that extend into water are called ice shelves; they tend to be thin with limited slopes and reduced velocities.[7] Narrow, fast-moving sections of an ice sheet are called ice streams.[8][9] In Antarctica, many ice streams drain into large ice shelves. Some drain directly into the sea, often with an ice tongue, like Mertz Glacier.

Sightseeing boat in front of a tidewater glacier, Kenai Fjords National Park, Alaska

Tidewater glaciers are glaciers that terminate in the sea, including most glaciers flowing from Greenland, Antarctica, Baffin and Ellesmere Islands in Canada, Southeast Alaska, and the Northern and Southern Patagonian Ice Fields. As the ice reaches the sea, pieces break off, or calve, forming icebergs. Most tidewater glaciers calve above sea level, which often results in a tremendous impact as the iceberg strikes the water. Tidewater glaciers undergo centuries-long cycles of advance and retreat that are much less affected by the climate change than those of other glaciers.

Thermally, a temperate glacier is at melting point throughout the year, from its surface to its base. The ice of a polar glacier is always below freezing point from the surface to its base, although the surface snowpack may experience seasonal melting. A sub-polar glacierincludes both temperate and polar ice, depending on depth beneath the surface and position along the length of the glacier. In a similar way, the thermal regime of a glacier is often described by the temperature at its base alone. A cold-based glacier is below freezing at the ice-ground interface, and is thus frozen to the underlying substrate. A warm-based glacier is above or at freezing at the interface, and is able to slide at this contact.[10] This contrast is thought to a large extent to govern the ability of a glacier to effectively erode its bed, as sliding ice promotes plucking at rock from the surface below.[11] Glaciers which are partly cold-based and partly warm-based are known as polythermal.[10]

Formation

Gorner Glacier in Switzerland

Glaciers form where the accumulation of snow and ice exceeds ablation. The area in which a glacier forms is called a cirque (corrie or cwm) – a typically armchair-shaped geological feature (such as a depression between mountains enclosed by arêtes) – which collects and compresses through gravity the snow which falls into it. This snow collects and is compacted by the weight of the snow falling above it forming névé. Further crushing of the individual snowflakes and squeezing the air from the snow turns it into extremely dense ‘glacial ice’. This glacial ice will fill the cirque until it ‘overflows’ through a geological weakness or vacancy, such as the gap between two mountains. When the mass of snow and ice is sufficiently thick, it begins to move due to a combination of surface slope, gravity and pressure. On steeper slopes, this can occur with as little as 15 m (50 ft) of snow-ice.

In temperate glaciers, snow repeatedly freezes and thaws, changing into granular ice called firn. Under the pressure of the layers of ice and snow above it, this granular ice fuses into denser and denser firn. Over a period of years, layers of firn undergo further compaction and become glacial ice. Glacier ice is slightly less dense than ice formed from frozen water because it contains tiny trapped air bubbles.

Glacial ice has a distinctive blue tint because it absorbs some red light due to an overtone of the infrared OH stretching mode of the water molecule. Liquid water is blue for the same reason. The blue of glacier ice is sometimes misattributed to Rayleigh scattering due to bubbles in the ice.[12]

A glacier cave located on the Perito Moreno Glacier in Argentina.

Structure

A glacier originates at a location called its glacier head and terminates at its glacier foot, or terminus.

Glaciers are broken into zones based on surface snowpack and melt conditions.[13] The ablation zone is the region where there is a net loss in glacier mass. The equilibrium line separates the ablation zone and the accumulation zone; it is the altitude where the amount of new snow gained by accumulation is equal to the amount of ice lost through ablation. The upper part of a glacier, where accumulation exceeds ablation, is called the accumulation zone. In general, the accumulation zone accounts for 60–70% of the glacier’s surface area, more if the glacier calves icebergs. Ice in the accumulation zone is deep enough to exert a downward force that erodes underlying rock. After a glacier melts, it often leaves behind a bowl- or amphitheater-shaped depression that ranges in size from large basins like the Great Lakes to smaller mountain depressions known as cirques.

The accumulation zone can be subdivided based on its melt conditions.

  1. The dry snow zone is a region where no melt occurs, even in the summer, and the snowpack remains dry.
  2. The percolation zone is an area with some surface melt, causing meltwater to percolate into the snowpack. This zone is often marked by refrozen ice lenses, glands, and layers. The snowpack also never reaches melting point.
  3. Near the equilibrium line on some glaciers, a superimposed ice zone develops. This zone is where meltwater refreezes as a cold layer in the glacier, forming a continuous mass of ice.
  4. The wet snow zone is the region where all of the snow deposited since the end of the previous summer has been raised to 0 °C.

The “health” of a glacier is usually assessed by determining the glacier mass balance or observing terminus behavior. Healthy glaciers have large accumulation zones, more than 60% of their area snowcovered at the end of the melt season, and a terminus with vigorous flow.

Following the Little Ice Age’s end around 1850, glaciers around the Earth have retreated substantially. A slight cooling led to the advance of many alpine glaciers between 1950–1985, but since 1985 glacier retreat and mass loss has become larger and increasingly ubiquitous.[14][15][16]

Motion

Shear or herring-bone crevasses on Emmons Glacier (Mount Rainier); such crevasses often form near the edge of a glacier where interactions with underlying or marginal rock impede flow. In this case, the impediment appears to be some distance from the near margin of the glacier.

Main article: Ice sheet dynamics

Glaciers move, or flow, downhill due to gravity and the internal deformation of ice.[17] Ice behaves like a brittle solid until its thickness exceeds about 50 m (160 ft). The pressure on ice deeper than 50 m causes plastic flow. At the molecular level, ice consists of stacked layers of molecules with relatively weak bonds between layers. When the stress on the layer above exceeds the inter-layer binding strength, it moves faster than the layer below.[18]

Glaciers also move through basal sliding. In this process, a glacier slides over the terrain on which it sits, lubricated by the presence of liquid water. The water is created from ice that melts under high pressure from frictional heating. Basal sliding is dominant in temperate, or warm-based glaciers.

Fracture zone and cracks

Ice cracks in the Titlis Glacier

The top 50 metres (160 ft) of a glacier are rigid because they are under low pressure. This upper section is known as the fracture zone; it mostly moves as a single unit over the plastically flowing lower section. When a glacier moves through irregular terrain, cracks called crevasses develop in the fracture zone. Crevasses form due to differences in glacier velocity. If two rigid sections of a glacier move at different speeds and directions, shear forces cause them to break apart, opening a crevasse. Crevasses are seldom more than 150 feet (46 m) deep but in some cases can be 1,000 feet (300 m) or even deeper. Beneath this point, the plasticity of the ice is too great for cracks to form. Intersecting crevasses can create isolated peaks in the ice, called seracs.

Crevasses can form in several different ways. Transverse crevasses are transverse to flow and form where steeper slopes cause a glacier to accelerate. Longitudinal crevasses form semi-parallel to flow where a glacier expands laterally. Marginal crevasses form from the edge of the glacier, due to the reduction in speed caused by friction of the valley walls. Marginal crevasses are usually largely transverse to flow. Moving glacier ice can sometimes separate from stagnant ice above, forming a bergschrund. Bergschrunds resemble crevasses but are singular features at a glacier’s margins.

Crevasses make travel over glaciers hazardous, especially when they are hidden by fragile snow bridges.

Crossing a crevasse on the Easton Glacier, Mount Baker, in the North Cascades, United States

Below the equilibrium line, glacial meltwater is concentrated in stream channels. Meltwater can pool in proglacial lakes on top of a glacier or descend into the depths of a glacier via moulins. Streams within or beneath a glacier flow in englacial or sub-glacial tunnels. These tunnels sometimes reemerge at the glacier’s surface.[19]

Speed

The speed of glacial displacement is partly determined by friction. Friction makes the ice at the bottom of the glacier move more slowly than ice at the top. In alpine glaciers, friction is also generated at the valley’s side walls, which slows the edges relative to the center.

Mean speeds vary greatly, but is typically around 1 meter per day.[20] There may be no motion in stagnant areas; for example, in parts of Alaska, trees can establish themselves on surface sediment deposits. In other cases, glaciers can move as fast as 20–30 m per day, such as in Greenland’s Jakobshavn Isbræ (Greenlandic: Sermeq Kujalleq). Velocity increases with increasing slope, increasing thickness, increasing snowfall, increasing longitudinal confinement, increasing basal temperature, increasing meltwater production and reduced bed hardness.

A few glaciers have periods of very rapid advancement called surges. These glaciers exhibit normal movement until suddenly they accelerate, then return to their previous state. During these surges, the glacier may reach velocities far greater than normal speed.[21] These surges may be caused by failure of the underlying bedrock, the pooling of meltwater at the base of the glacier[22] — perhaps delivered from a supraglacial lake — or the simple accumulation of mass beyond a critical “tipping point”.[23]

In glaciated areas where the glacier moves faster than one km per year, glacial earthquakes occur. These are large scale tremblors that have seismic magnitudes as high as 6.1.[24][25] The number of glacial earthquakes in Greenland peaks every year in July, August and September and is increasing over time. In a study using data from January 1993 through October 2005, more events were detected every year since 2002, and twice as many events were recorded in 2005 as there were in any other year. This increase in the numbers of glacial earthquakes in Greenland may be a response to global warming.[24][25]

Ogives

Ogives are alternating wave crests and valleys that appear as dark and light bands of ice on glacier surfaces. They are linked to seasonal motion of glaciers; the width of one dark and one light band generally equals the annual movement of the glacier. Ogives are formed when ice from an icefall is severely broken up, increasing ablation surface area during summer. This creates a swale and space for snow accumulation in the winter, which in turn creates a ridge.[26] Sometimes ogives consist only of undulations or color bands and are described as wave ogives or band ogives.[27]

Geography

For more details on this topic, see List of glaciers, and Retreat of glaciers since 1850.

Black ice glacier near Aconcagua, Argentina

Glaciers are present on every continent and approximately fifty countries, excluding those (Australia, South Africa) that have glaciers only on distant subantarctic island territories. Extensive glaciers are found in Antarctica, Chile, Canada, Alaska, Greenland and Iceland. Mountain glaciers are widespread, especially in the Andes, the Himalayas, the Rocky Mountains, the Caucasus, and the Alps. Mainland Australia currently contains no glaciers, although a small glacier on Mount Kosciuszko was present in the last glacial period.[28] In New Guinea, small, rapidly diminishing, glaciers are located on its highest summit massif of Puncak Jaya.[29] Africa has glaciers on Mount Kilimanjaro in Tanzania, on Mount Kenya and in the Rwenzori Mountains. Oceanic islands with glaciers occur on Iceland, Svalbard, New Zealand, Jan Mayen and the subantarctic islands of Marion, Heard, Grande Terre (Kerguelen) and Bouvet. During glacial periods of the Quaternary, Taiwan, Hawaii on Mauna Kea[30] and Tenerife also had large alpine glaciers, while the Faroe and Crozet Islands[31] were completely glaciated.

The permanent snow cover necessary for glacier formation is affected by factors such as the degree of slope on the land, amount of snowfall and the winds. Glaciers can be found in all latitudes except from 20° to 27° north and south of the equator where the presence of the descending limb of the Hadley circulation lowers precipitation so much that with high insolation snow lines reach above 6,500 metres (21,330 ft). Between 19˚N and 19˚S, however, precipitation is higher and the mountains above 5,000 metres (16,400 ft) usually have permanent snow.

Even at high latitudes, glacier formation is not inevitable. Areas of the Arctic, such as Banks Island, and the McMurdo Dry Valleys in Antarctica are considered polar deserts where glaciers cannot form because they receive little snowfall despite the bitter cold. Cold air, unlike warm air, is unable to transport much water vapor. Even during glacial periods of the Quaternary, Manchuria, lowland Siberia,[32]and central and northern Alaska,[33] though extraordinarily cold, had such light snowfall that glaciers could not form.[34][35]

In addition to the dry, unglaciated polar regions, some mountains and volcanoes in Bolivia, Chile and Argentina are high (4,500 metres (14,800 ft) – 6,900 m (22,600 ft)) and cold, but the relative lack of precipitation prevents snow from accumulating into glaciers. This is because these peaks are located near or in the hyperarid Atacama Desert.

Glacial geology

Diagram of glacial plucking and abrasion

Glacially plucked granitic bedrock near Mariehamn, Åland Islands

Glaciers erode terrain through two principal processes: abrasion and plucking.

As glaciers flow over bedrock, they soften and lift blocks of rock into the ice. This process, called plucking, is caused by subglacial water that penetrates fractures in the bedrock and subsequently freezes and expands. This expansion causes the ice to act as a lever that loosens the rock by lifting it. Thus, sediments of all sizes become part of the glacier’s load. If a retreating glacier gains enough debris, it may become a rock glacier, like the Timpanogos Glacier in Utah.

Abrasion occurs when the ice and its load of rock fragments slide over bedrock and function as sandpaper, smoothing and polishing the bedrock below. The pulverized rock this process produces is called rock flour and is made up of rock grains between 0.002 and 0.00625 mm in size. Abrasion leads to steeper valley walls and mountain slopes in alpine settings, which can cause avalanches and rock slides. These add even more material to the glacier.

Glacial abrasion is commonly characterized by glacial striations. Glaciers produce these when they contain large boulders that carve long scratches in the bedrock. By mapping the direction of the striations, researchers can determine the direction of the glacier’s movement. Similar to striations are chatter marks, lines of crescent-shape depressions in the rock underlying a glacier. They are formed by abrasion when boulders in the glacier are repeatedly caught and released as they are dragged along the bedrock.

The rate of glacier erosion is variable. Six factors control erosion rate:

  • Velocity of glacial movement
  • Thickness of the ice
  • Shape, abundance and hardness of rock fragments contained in the ice at the bottom of the glacier
  • Relative ease of erosion of the surface under the glacier
  • Thermal conditions at the glacier base
  • Permeability and water pressure at the glacier base

Material that becomes incorporated in a glacier is typically carried as far as the zone of ablation before being deposited. Glacial deposits are of two distinct types:

  • Glacial till: material directly deposited from glacial ice. Till includes a mixture of undifferentiated material ranging from clay size to boulders, the usual composition of a moraine.
  • Fluvial and outwash sediments: sediments deposited by water. These deposits are stratified by size.

Larger pieces of rock that are encrusted in till or deposited on the surface are called “glacial erratics”. They range in size from pebbles to boulders, but as they are often moved great distances, they may be drastically different from the material upon which they are found. Patterns of glacial erratics hint at past glacial motions.

Moraines

Glacial moraines above Lake Louise, Alberta, Canada

Glacial moraines are formed by the deposition of material from a glacier and are exposed after the glacier has retreated. They usually appear as linear mounds of till, a non-sorted mixture of rock, gravel and boulders within a matrix of a fine powdery material. Terminal or end moraines are formed at the foot or terminal end of a glacier. Lateral moraines are formed on the sides of the glacier. Medial moraines are formed when two different glaciers merge and the lateral moraines of each coalesce to form a moraine in the middle of the combined glacier. Less apparent are ground moraines, also called glacial drift, which often blankets the surface underneath the glacier downslope from the equilibrium line.

The term moraine is of French origin. It was coined by peasants to describe alluvial embankments and rims found near the margins of glaciers in the French Alps. In modern geology, the term is used more broadly, and is applied to a series of formations, all of which are composed of till. Moraines can also create moraine dammed lakes.

Drumlins

A drumlin field forms after a glacier has modified the landscape. The teardrop-shaped formations denote the direction of the ice flow.

Drumlins are asymmetrical, canoe shaped hills made mainly of till. Their heights vary from 15 to 50 meters and they can reach a kilometer in length. The tilted side of the hill faces the direction from which the ice advanced (stoss), while the longer slope follows the ice’s direction of movement (lee).

Drumlins are found in groups called drumlin fields or drumlin camps. One of these fields is found east of Rochester, New York; it is estimated to contain about 10,000 drumlins.

Although the process that forms drumlins is not fully understood, their shape implies that they are products of the plastic deformation zone of ancient glaciers. It is believed that many drumlins were formed when glaciers advanced over and altered the deposits of earlier glaciers.

Glacial valleys, cirques, arêtes, and pyramidal peaks

Features of a glacial landscape

Before glaciation, mountain valleys have a characteristic “V” shape, produced by eroding water. During glaciation, these valleys are widened, deepened, and smoothed, forming a “U”-shaped glacial valley. The erosion that creates glacial valleys eliminates the spurs of earth that extend across mountain valleys, creating triangular cliffs called truncated spurs. Within glacial valleys, depressions created by plucking and abrasion can be filled by lakes, called paternoster lakes. If a glacial valley runs into a large body of water, it forms a fjord.

Many glaciers deepen their valleys more than their smaller tributaries. Therefore, when glaciers recede, the valleys of the tributary glaciers remain above the main glacier’s depression and are called hanging valleys.

At the start of a classic valley glacier is a bowl-shaped cirque, which has escarped walls on three sides but is open on the side that descends into the valley. Cirques are where ice begins to accumulate in a glacier. Two glacial cirques may form back to back and erode their backwalls until only a narrow ridge, called an arête is left. This structure may result in a mountain pass. If multiple cirques encircle a single mountain, they create pointed pyramidal peaks; particularly steep examples are called horns.

Roche moutonnée

Some rock formations in the path of a glacier are sculpted into small hills called roche moutonnée, or “sheepback” rock. Roche moutonnée are elongated, rounded, and asymmetrical bedrock knobs can be produced by glacier erosion. They range in length from less than a meter to several hundred meters long.[36] Roche moutonnée have a gentle slope on their up-glacier sides and a steep to vertical face on their down-glacier sides. The glacier abrades the smooth slope on the upstream side as it flows along, but tears loose and carries away rock from the downstream side via plucking.

Alluvial stratification

As the water that rises from the ablation zone moves away from the glacier, it carries fine eroded sediments with it. As the speed of the water decreases, so does its capacity to carry objects in suspension. The water thus gradually deposits the sediment as it runs, creating an alluvial plain. When this phenomenon occurs in a valley, it is called a valley train. When the deposition is in an estuary, the sediments are known as bay mud.

Outwash plains and valley trains are usually accompanied by basins known as “kettles”. These are small lakes formed when large ice blocks that are trapped in alluvium melt and produce water-filled depressions. Kettle diameters range from 5 m to 13 km, with depths of up to 45 meters. Most are circular in shape because the blocks of ice that formed them were rounded as they melted.[37]

Glacial deposits

Landscape produced by a receding glacier

When a glacier’s size shrinks below a critical point, its flow stops and it becomes stationary. Meanwhile, meltwater within and beneath the ice leaves stratified alluvial deposits. These deposits, in the forms of columns, terraces and clusters, remain after the glacier melts and are known as “glacial deposits”.

Glacial deposits that take the shape of hills or mounds are called kames. Some kames form when meltwater deposits sediments through openings in the interior of the ice. Others are produced by fans or deltas created by meltwater. When the glacial ice occupies a valley, it can form terraces or kames along the sides of the valley.

Long, sinuous glacial deposits are called eskers. Eskers are composed of sand and gravel that was deposited by meltwater streams that flowed through ice tunnels within or beneath a glacier. They remain after the ice melts, with heights exceeding 100 meters and lengths of as long as 100 km.

Loess deposits

Very fine glacial sediments or rock flour is often picked up by wind blowing over the bare surface and may be deposited great distances from the original fluvial deposition site. These eolian loess deposits may be very deep, even hundreds of meters, as in areas of China and the Midwestern United States of America. Katabatic winds can be important in this process.

Isostatic rebound

Main article: Isostatic rebound

Isostatic pressure by a glacier on the Earth’s crust

Large masses, such as ice sheets or glaciers, can depress the crust of the Earth into the mantle. The depression usually totals a third of the ice sheet or glacier’s thickness. After the ice sheet or glacier melts, the mantle begins to flow back to its original position, pushing the crust back up. This post-glacial rebound, which proceeds very slowly after the melting of the ice sheet or glacier, is currently occurring in measurable amounts in Scandinavia and the Great Lakes region of North America.

A geomorphological feature created by the same process on a smaller scale is known as dilation-faulting. It occurs where previously compressed rock is allowed to return to its original shape more rapidly than can be maintained without faulting. This leads to an effect similar to what would be seen if the rock were hit by a large hammer. Dilation faulting can be observed in recently de-glaciated parts of Iceland and Cumbria.

On Mars

Northern polar ice cap on Mars

Main article: Glaciers on Mars

The polar ice caps of Mars show geologic evidence of glacial deposits. The south polar cap is especially comparable to glaciers on Earth.[38] Topographical features and computer models indicate the existence of more glaciers in Mars’ past.[39]

At mid-latitudes, between 35° and 65° north or south, Martian glaciers are affected by the thin Martian atmosphere. Because of the low atmospheric pressure, ablation near the surface is solely due to sublimation, not melting. As on Earth, many glaciers are covered with a layer of rocks which insulates the ice. A radar instrument on board the Mars Reconnaissance Orbiter found ice under a thin layer of rocks in formations called Lobate Debris Aprons (LDAs).[40][41][42][43][44]

Notes

  1. ^ Post, Austin; LaChapelle, Edward R (2000). Glacier ice. Seattle, Washington: University of Washington Press. ISBN 0-295-97910-0.
  2. ^ Brown, Molly Elizabeth; Ouyang, Hua; Habib, Shahid; Shrestha, Basanta; Shrestha, Mandira; Panday, Prajjwal; Tzortziou, Maria; Policelli, Frederick; Artan, Guleid; Giriraj, Amarnath; Bajracharya, Sagar R.; Racoviteanu, Adina. “HIMALA: Climate Impacts on Glaciers, Snow, and Hydrology in the Himalayan Region”. Mountain Research and Development. International Mountain Society. Retrieved 16 September 2011.
  3. ^ Simpson, D.P. (1979). Cassell’s Latin Dictionary (5 ed.). London: Cassell Ltd. p. 883. ISBN 0-304-52257-0.
  4. ^ “Retreat of Alaskan glacier Juneau icefield”. Nichols.edu. Retrieved 2009-01-05.
  5. ^ “American Meteorological Society, Glossary of Meteorology”. Amsglossary.allenpress.com. Retrieved 2013-01-04.
  6. ^ “Sea Level and Climate”. USGS FS 002-00. USGS. 2000-01-31. Retrieved 2009-01-05.
  7. ^ * National Snow and Ice Data Center. “Types of Glacier”.
  8. ^ Bindschadler, R.A. and T.A. Scambos. Satellite-image-derived velocity field of an Antarctic ice stream. Science, 252(5003), 242-246, 1991
  9. ^ British Antarctic Survey. “Description of Ice Streams”. Retrieved 2009-01-26.
  10. ^ a b http://link.springer.com/referenceworkentry/10.1007%2F978-90-481-2642-2_72/fulltext.html
  11. ^ Boulton, G.S. [1974] “Processes and patterns of glacial erosion”, (In Coates, D.R. ed., Glacial Geomorphology. A Proceedings Volume of the fifth Annual Geomorphology Symposia series, held at Binghamton, New York, September 26–28, 1974. Binghamton, N.Y., State University of New York, p. 41-87. (Publications in Geomorphology))
  12. ^ “What causes the blue color that sometimes appears in snow and ice ?”. Webexhibits.org. Retrieved 2013-01-04.
  13. ^ Benson, C.S., 1961, “Stratigraphic studies in the snow and firn of the Greenland Ice Sheet”, Res. Rep. 70, U.S. Army Snow, Ice and Permafrost Res Establ., Corps of Eng., 120 pp
  14. ^ “Glacier change and related hazards in Switzerland”. UNEP. Retrieved 2009-01-05.
  15. ^ “Frank Paul, et al., 2004, Rapid disintegration of Alpine glaciers observed with satellite data, GEOPHYSICAL RESEARCH LETTERS, VOL. 31, L21402,

    ParseError: EOF expected (click for details)

    Callstack:at (Courses/Lumen_Learning/Book:_Earth_Science_(Lumen)/21:_Glaciers/21.2:_Glaciers), /content/body/div[1]/div/div[14]/div/ol/li[15]/a/span, line 1, column 4
    ” (PDF). 2004.
  16. ^ “Recent Global Glacier Retreat Overview” (PDF). Retrieved 2013-01-04.
  17. ^ Greve, R.; Blatter, H. (2009). Dynamics of Ice Sheets and Glaciers. Springer. doi:10.1007/978-3-642-03415-2. ISBN 978-3-642-03414-5.
  18. ^ W.S.B. Paterson, Physics of ice
  19. ^ “Moulin ‘Blanc’: NASA Expedition Probes Deep Within a Greenland Glacier”. NASA. 2006-12-11. Retrieved 2009-01-05.
  20. ^ Glacier properties Hunter College CUNY lectures
  21. ^ T. Strozzi et al.: The Evolution of a Glacier Surge Observed with the ERS Satellites (pdf, 1.3 Mb)
  22. ^ “The Brúarjökull Project: Sedimentary environments of a surging glacier. The Brúarjökull Project research idea”. Hi.is. Retrieved 2013-01-04.
  23. ^ Meier & Post (1969)
  24. ^ a b http://people.deas.harvard.edu/~vtsai/files/EkstromNettlesTsai_Science2006.pdf Ekström, G., M. Nettles, and V. C. Tsai (2006)”Seasonality and Increasing Frequency of Greenland Glacial Earthquakes”, Science, 311, 5768, 1756-1758, doi:10.1126/science.1122112
  25. ^ a b http://people.deas.harvard.edu/~vtsai/files/TsaiEkstrom_JGR2007.pdf Tsai, V. and G. Ekström (2007). “Analysis of Glacial Earthquakes”, J. Geophys. Res., 112, F03S22, doi:10.1029/2006JF000596
  26. ^ Easterbrook, D.J. (1999). Surface Processes and Landforms (2 ed.). New Jersey: Prentice-Hall, Inc. p. 546. ISBN 0-13-860958-6.
  27. ^ “Glossary of Glacier Terminology”. Pubs.usgs.gov. 2012-06-20. Retrieved 2013-01-04.
  28. ^ “C.D. Ollier: ”Australian Landforms and their History”, National Mapping Fab, Geoscience Australia”. Ga.gov.au. 2010-11-18. Retrieved 2013-01-04.
  29. ^ KINCAID, JONI L.; KLEIN, ANDREW G. (2004). “Retreat of the Irian Jaya Glaciers from 2000 to 2002 as Measured from IKONOS Satellite Images”. Portland, Maine, USA. pp. 147–157. Retrieved 2009-01-05.
  30. ^ “Hawaiian Glaciers Reveal Clues to Global Climate Change”. Geology.com. 2007-01-26. Retrieved 2013-01-04.
  31. ^ “French Colonies – Crozet Archipelago”. Discoverfrance.net. 2010-12-09. Retrieved 2013-01-04.
  32. ^ Collins, Henry Hill; Europe and the USSR; p. 263. ISBN 1-256-35000-3
  33. ^ “Yukon Beringia Interpretive Center”. Beringia.com. 1999-04-12. Retrieved 2013-01-04.
  34. ^ Earth History 2001 (page 15)
  35. ^ “On the Zoogeography of the Holarctic Region”. Wku.edu. Retrieved 2013-01-04.
  36. ^ ‘Glaciers & Glaciation’ (Arnold, London 1998) Douglas Benn and David Evans, pp324-326
  37. ^ “Kettle geology”. Britannica Online. Retrieved 2009-03-12.
  38. ^ “Kargel, J.S. et al.:”Martian Polar Ice Sheets and Mid-Latitude Debris-Rich Glaciers, and Terrestrial Analogs”, Third International Conference on Mars Polar Science and Exploration, Alberta, Canada, October 13-17, 2003 (pdf 970 Kb)” (PDF). Retrieved 2013-01-04.
  39. ^ “Martian glaciers: did they originate from the atmosphere? ESA Mars Express, 20 January 2006”. Esa.int. 2006-01-20. Retrieved 2013-01-04.
  40. ^ Head, J. et al. 2005. Tropical to mid-latitude snow and ice accumulation, flow and glaciation on Mars. Nature: 434. 346-350
  41. ^ Source: Brown University Posted Monday, October 17, 2005 (2005-10-17). “Mars’ climate in flux: Mid-latitude glaciers | SpaceRef – Your Space Reference”. Marstoday.com. Retrieved 2013-01-04.
  42. ^ Richard Lewis (2008-04-23). “Glaciers Reveal Martian Climate Has Been Recently Active | Brown University News and Events”. News.brown.edu. Retrieved 2013-01-04.
  43. ^ Plaut, J. 2008. Radar Evidence for Ice in Lobate Debris Aprons in the Mid-Northern Latitudes of Mars. Lunar and Planetary Science XXXIX. 2290.pdf
  44. ^ Holt, J. Radar Sounding Evidence for Ice within Lobate Debris Aprons near Hellas Basin, Mid-Southern Latitudes of Mars. 2441.pdf

References

  • This article draws heavily on the corresponding article in the Spanish-language Wikipedia, which was accessed in the version of 24 July 2005.
  • Hambrey, Michael; Alean, Jürg (2004). Glaciers (2nd ed.). Cambridge University Press. ISBN 0-521-82808-2. OCLC 54371738. An excellent less-technical treatment of all aspects, with superb photographs and firsthand accounts of glaciologists’ experiences. All images of this book can be found online (see Weblinks: Glaciers-online)
  • Benn, Douglas I.; Evans, David J. A. (1999). Glaciers and Glaciation. Arnold. ISBN 0-470-23651-5. OCLC 38329570.
  • Bennett, M. R.; Glasser, N. F. (1996). Glacial Geology: Ice Sheets and Landforms. John Wiley & Sons. ISBN 0-471-96344-5. OCLC 33359888 37536152.
  • Hambrey, Michael (1994). Glacial Environments. University of British Columbia Press, UCL Press. ISBN 0-7748-0510-2. OCLC 30512475. An undergraduate-level textbook.
  • Knight, Peter G (1999). Glaciers. Cheltenham: Nelson Thornes. ISBN 0-7487-4000-7. OCLC 42656957 63064183 77294832. A textbook for undergraduates avoiding mathematical complexities
  • Walley, Robert (1992). Introduction to Physical Geography. Wm. Brown Publishers. A textbook devoted to explaining the geography of our planet.
  • W. S. B. Paterson (1994). Physics of Glaciers (3rd ed.). Pergamon Press. ISBN 0-08-013972-8. OCLC 26188. A comprehensive reference on the physical principles underlying formation and behavior.

21.2: Glaciers - Geosciences

Laurentia extends as far west as eastern B.C. (Figure 21.3), but the ancient rocks of the craton are almost completely covered by younger rocks in B.C., Yukon, and all of Alberta except the far northeast corner. Laurentia is well represented in northern Saskatchewan and across large parts of Manitoba, the Northwest Territories, and Nunavut (Figure 21.5). Where they are exposed, the rocks of the Canadian Shield are highly varied lithologically, typically strongly metamorphosed due to their deep burial at some time in the past, and in some cases, quite different from what could be expected to occur on Earth today.

Starting from the south, in eastern Manitoba and adjacent Ontario, we have the ancient rocks of the Superior Province. On the map the Superior Province, rocks are mostly pink, representing granitic and gneissic rocks, with strips and blotches of green, representing metamorphosed sea-floor basalt and sediments, also known as greenstone belts. These rocks are widely interpreted to have deep crustal origins, and include large areas of granulite facies metamorphic rock formed at high temperatures and moderate to high pressures (see Figure 7.19). Superior Province greenstone belts in Ontario and Quebec host some of the world’s largest volcanogenic massive sulphide deposits. As described in Chapter 20, the Superior Province in northern Manitoba is host to important nickel deposits at Thompson. These formed from mantle-derived mafic magma that interacted with sulphur-bearing crustal rocks, and within which heavy-metal sulphide minerals formed.

The Trans-Hudson Orogen (THO), as its name implies, extends through Saskatchewan and Manitoba and over to the eastern side of Hudson Bay. It represents the continent-continent collision zone between the Superior Craton to the south and the Churchill Craton (including the Wyoming, Hearne, and Rae Cratons) to the north thus it’s a remnant of the initial formation of Laurentia at around 1.9 Ga. At the time of the collision, the THO would have been a major mountain range, and the rocks that we see there now — which evolved deep beneath those mountains — are highly metamorphosed sedimentary and volcanic rocks intruded by large granitic bodies. The important volcanogenic massive sulphide deposits around Flin Flon are within the THO.

Figure 21.5 Geological features of the Canadian Shield of western Canada. A.B.: Athabasca Basin, T.B.: Thelon Basin, and TMZ: Taltson Magmatic Zone [ By SE after: http://geoscan.nrcan.gc.ca/starweb/geoscan/servlet.starweb?path=geoscan/fulle.web&search1=R=208175]

The Churchill Craton is lithologically similar to the Superior Craton, although not generally as old. It includes two important sedimentary basins: the Athabasca Basin in Saskatchewan and the Thelon Basin in Nunavut, both filled with rocks aged around 1.7 Ga. These consist primarily of sandstones and minor mudstones that are only weakly metamorphosed and essentially undeformed (not folded) because they are situated within a stable craton and so have not been subjected to significant tectonic forces. The Athabasca Basin is economically important for its large and rich unconformity-type uranium deposits (see Chapter 20). At its western end, there is the remnant of a large extraterrestrial impact, the 40 km diameter Carswell Crater. When the meteor struck at this location, at around 115 Ma, the impact and subsequent rebound of the crust was enough to bring metamorphic rock up to surface from beneath about 2,000 m of Athabasca Group sandstone. There is no connection between the Carswell Crater and the much older (

The Taltson Magmatic Zone (TMZ), which forms the boundary between the Churchill and Slave Cratons, consists primarily of granitic rock. One interpretation is that the TMZ formed along a convergent boundary, although this is not universally accepted.

The Slave Craton is dominated by granitic rocks and metamorphosed clastic sedimentary rocks. On its western edge, there is a large area of very old gneissic rock that includes the Acasta Gneiss, dated at 4.03 Ga, which, for the time being at least, is the oldest rock in the world (Figure 21.6).

Figure 21.6 A sample of the Acasta Gneiss on display at the Natural History Museum in Vienna [https://commons.wikimedia.org/wiki/File:Acasta_gneiss.jpg]

The Wopmay Orogen, interpreted as the site of another ancient continent-continent collision, lies to the west of the Slave Craton. Although mostly composed of felsic igneous rocks and gneisses, the Wopmay Orogen includes a body of mafic and ultramafic igneous rock called the Muskox Intrusion. Derived from a mantle plume and dated at about 1.1 Ga, the Muskox is comparable to a handful of other mafic and ultramafic intrusions around the world in that it has distinctive repetitive layering caused by settling of heavy metal-rich minerals within the low-viscosity magma. Muskox has high levels of nickel, copper, and chromium, and has the potential to have platinum and palladium like a similar body in South Africa. Ultramafic intrusions like Muskox do not take place on Earth today because the mantle is no longer hot enough.

The oldest rocks in British Columbia are the strongly metamorphosed sedimentary, volcanic, and intrusive rocks of the Monashee Complex, situated to the west of the Columbia River near Revelstoke (Figure 21.7). Aged around 2 Ga, these may actually be part of Laurentia.


138 21.2 Western Canada during the Precambrian

Laurentia extends as far west as eastern B.C. (Figure 21.3), but the ancient rocks of the craton are almost completely covered by younger rocks in B.C., Yukon, and all of Alberta except the far northeast corner. Laurentia is well represented in northern Saskatchewan and across large parts of Manitoba, the Northwest Territories, and Nunavut (Figure 21.5). Where they are exposed, the rocks of the Canadian Shield are highly varied lithologically, typically strongly metamorphosed due to their deep burial at some time in the past, and in some cases, quite different from what could be expected to occur on Earth today.

Starting from the south, in eastern Manitoba and adjacent Ontario, we have the ancient rocks of the Superior Province. On the map the Superior Province, rocks are mostly pink, representing granitic and gneissic rocks, with strips and blotches of green, representing metamorphosed sea-floor basalt and sediments, also known as greenstone belts. These rocks are widely interpreted to have deep crustal origins, and include large areas of granulite facies metamorphic rock formed at high temperatures and moderate to high pressures (see Figure 7.19). Superior Province greenstone belts in Ontario and Quebec host some of the world’s largest volcanogenic massive sulphide deposits. As described in Chapter 20, the Superior Province in northern Manitoba is host to important nickel deposits at Thompson. These formed from mantle-derived mafic magma that interacted with sulphur-bearing crustal rocks, and within which heavy-metal sulphide minerals formed.

The Trans-Hudson Orogen (THO), as its name implies, extends through Saskatchewan and Manitoba and over to the eastern side of Hudson Bay. It represents the continent-continent collision zone between the Superior Craton to the south and the Churchill Craton (including the Wyoming, Hearne, and Rae Cratons) to the north thus it’s a remnant of the initial formation of Laurentia at around 1.9 Ga. At the time of the collision, the THO would have been a major mountain range, and the rocks that we see there now — which evolved deep beneath those mountains — are highly metamorphosed sedimentary and volcanic rocks intruded by large granitic bodies. The important volcanogenic massive sulphide deposits around Flin Flon are within the THO.

Figure 21.5 Geological features of the Canadian Shield of western Canada. A.B.: Athabasca Basin, T.B.: Thelon Basin, and TMZ: Taltson Magmatic Zone [ By SE after: http://geoscan.nrcan.gc.ca/starweb/geoscan/servlet.starweb?path=geoscan/fulle.web&search1=R=208175]

The Churchill Craton is lithologically similar to the Superior Craton, although not generally as old. It includes two important sedimentary basins: the Athabasca Basin in Saskatchewan and the Thelon Basin in Nunavut, both filled with rocks aged around 1.7 Ga. These consist primarily of sandstones and minor mudstones that are only weakly metamorphosed and essentially undeformed (not folded) because they are situated within a stable craton and so have not been subjected to significant tectonic forces. The Athabasca Basin is economically important for its large and rich unconformity-type uranium deposits (see Chapter 20). At its western end, there is the remnant of a large extraterrestrial impact, the 40 km diameter Carswell Crater. When the meteor struck at this location, at around 115 Ma, the impact and subsequent rebound of the crust was enough to bring metamorphic rock up to surface from beneath about 2,000 m of Athabasca Group sandstone. There is no connection between the Carswell Crater and the much older (

The Taltson Magmatic Zone (TMZ), which forms the boundary between the Churchill and Slave Cratons, consists primarily of granitic rock. One interpretation is that the TMZ formed along a convergent boundary, although this is not universally accepted.

The Slave Craton is dominated by granitic rocks and metamorphosed clastic sedimentary rocks. On its western edge, there is a large area of very old gneissic rock that includes the Acasta Gneiss, dated at 4.03 Ga, which, for the time being at least, is the oldest rock in the world (Figure 21.6).

Figure 21.6 A sample of the Acasta Gneiss on display at the Natural History Museum in Vienna [https://commons.wikimedia.org/wiki/File:Acasta_gneiss.jpg]

The Wopmay Orogen, interpreted as the site of another ancient continent-continent collision, lies to the west of the Slave Craton. Although mostly composed of felsic igneous rocks and gneisses, the Wopmay Orogen includes a body of mafic and ultramafic igneous rock called the Muskox Intrusion. Derived from a mantle plume and dated at about 1.1 Ga, the Muskox is comparable to a handful of other mafic and ultramafic intrusions around the world in that it has distinctive repetitive layering caused by settling of heavy metal-rich minerals within the low-viscosity magma. Muskox has high levels of nickel, copper, and chromium, and has the potential to have platinum and palladium like a similar body in South Africa. Ultramafic intrusions like Muskox do not take place on Earth today because the mantle is no longer hot enough.

The oldest rocks in British Columbia are the strongly metamorphosed sedimentary, volcanic, and intrusive rocks of the Monashee Complex, situated to the west of the Columbia River near Revelstoke (Figure 21.7). Aged around 2 Ga, these may actually be part of Laurentia.

Figure 21.7 Precambrian rocks in southern B.C. and Alberta [ By SE after: http://geoscan.nrcan.gc.ca/starweb/geoscan/servlet.starweb?path=geoscan/fulle.web&search1=R=208175]

There are much more extensive Precambrian rocks within the Columbia and Rocky Mountains of southeastern B.C. and the southwestern corner of Alberta. The rocks of the Purcell Supergroup (a supergroup comprises more than one group) are present in the extreme southeastern corner of B.C. and adjacent Alberta, and extend well into the United States (as the Belt Supergroup). These are mostly unmetamorphosed clastic rocks deposited in rivers and lakes during the middle Proterozoic, at around 1,400 Ma, while Laurentia was still part of the supercontinent Columbia. When Columbia rifted apart, the division happened within the area of the Purcell/Belt rocks. Similar rocks of the same age are present in Tasmania and Siberia, and it is postulated that they were once part of the same depositional basin.

Exercises

Exercise 21.2 Purcell Rocks Down Under?

This map shows the geology of the Australian state of Tasmania. Identify which rocks might be comparable to the Purcell rocks of B.C. and Alberta.

In what way are these rocks different from those in Canada?

The Windermere Group rocks — also mostly clastic sedimentary — were deposited in the ocean along the western edge of Laurentia (Figure 21.7) in the late Proterozoic (around 700 Ma) after the breakup of Columbia. In fact, sedimentary rocks of this age extend all along the western side of the Rocky Mountains, well into Yukon. Deposition in this area was taking place during the late Proterozoic Snowball Earth glaciations, as can be seen in Windermere Group rocks of the Toby Formation from the area south of Cranbrook, B.C. (Figure 21.8). The Toby Formation is a fine-grained marine rock (mudstone) with numerous large angular clasts of limestone and quartz. The mud was deposited in the quiet water of a continental slope environment, and the large clasts were dropped from floating ice derived from glaciers on Laurentia. The Toby Formation is unique in this area most of the rest of the late Proterozoic clastic sedimentary rocks in this region do not have glacial dropstones.

Figure 21.8 Late Proterozoic Toby Formation mudstone with glacial dropstones south of Cranbrook, B.C. [SE]


16.1 Glacial Periods in Earth’s History

We are currently in the middle of a glacial period (although it’s less intense now than it was 20,000 years ago) but this is not the only period of glaciation in Earth’s history there have been many in the distant past, as illustrated in Figure 16.2. In general, however, Earth has been warm enough to be ice-free for much more of the time than it has been cold enough to be glaciated.

Figure 16.2 The record of major past glaciations during Earth’s history. [SE]

The oldest known glacial period is the Huronian. Based on evidence of glacial deposits from the area around Lake Huron in Ontario and elsewhere, it is evident that the Huronian Glaciation lasted from approximately 2,400 to 2,100 Ma. Because rocks of that age are rare, we don’t know much about the intensity or global extent of this glaciation.

Late in the Proterozoic, for reasons that are not fully understood, the climate cooled dramatically and Earth was seized by what appears to be its most intense glaciation. The glaciations of the Cryogenian Period (cryo is Latin for icy cold) are also known as the “Snowball Earth” glaciations, because it is hypothesized that the entire planet was frozen — even in equatorial regions — with ice on the oceans up to 1 km thick. A visitor to our planet at that time might not have held out much hope for its inhabitability, although life still survived in the oceans. There were two main glacial periods within the Cryogenian, each lasting for about 20 million years: the Sturtian at around 700 Ma and the Marinoan at 650 Ma. There is also evidence of some shorter glaciations both before and after these. The end of the Cryogenian glaciations coincides with the evolution of relatively large and complex life forms on Earth. This started during the Ediacaran Period, and then continued with the so-called explosion of life forms in the Cambrian. Some geologists think that the changing environmental conditions of the Cryogenian are what actually triggered the evolution of large and complex life.

There have been three major glaciations during the Phanerozoic (the past 540 million years), including the Andean/Saharan (recorded in rocks of South America and Africa), the Karoo (named for rocks in southern Africa), and the Cenozoic glaciations. The Karoo was the longest of the Phanerozoic glaciations, persisting for much of the time that the supercontinent Gondwana was situated over the South Pole (

360 to 260 Ma). It covered large parts of Africa, South America, Australia, and Antarctica (see Figure 10.4). As you might recall from Chapter 10, this widespread glaciation, across continents that are now far apart, was an important component of Alfred Wegener’s evidence for continental drift. Unlike the Cryogenian glaciations, the Andean/Saharan, Karoo, and Cenozoic glaciations only affected parts of Earth. During Karoo times, for example, what is now North America was near the equator and remained unglaciated.

Earth was warm and essentially unglaciated throughout the Mesozoic. Although there may have been some alpine glaciation at this time, there is no longer any record of it. The dinosaurs, which dominated terrestrial habitats during the Mesozoic, did not have to endure icy conditions.

A warm climate persisted into the Cenozoic in fact there is evidence that the Paleocene (

50 to 60 Ma) was the warmest part of the Phanerozoic since the Cambrian (Figure 16.3). A number of tectonic events during the Cenozoic contributed to persistent and significant planetary cooling since 50 Ma. For example, the collision of India with Asia and the formation of the Himalayan range and the Tibetan Plateau resulted in a dramatic increase in the rate of weathering and erosion. Higher than normal rates of weathering of rocks with silicate minerals, especially feldspar, consumes carbon dioxide from the atmosphere and therefore reduces the greenhouse effect, resulting in long-term cooling.

Figure 16.3 The global temperature trend over the past 65 Ma (the Cenozoic). From the end of the Paleocene to the height of the Pleistocene Glaciation, global average temperature dropped by about 14°C. (PETM is the Palecene-Eocene thermal maximum) [SE after Routledge, 2013, http://www.alpineanalytics.com/Climate/DeepTime.html ]

At 40 Ma, ongoing plate motion widened the narrow gap between South America and Antarctica, resulting in the opening of the Drake Passage. This allowed for the unrestricted west-to-east flow of water around Antarctica, the Antarctic Circumpolar Current (Figure 16.4), which effectively isolated the southern ocean from the warmer waters of the Pacific, Atlantic, and Indian Oceans. The region cooled significantly, and by 35 Ma (Oligocene) glaciers had started to form on Antarctica.

Figure 16.4 The Antarctic Circumpolar Current (red arrows) prevents warm water from the rest of Earth’s oceans from getting close to Antarctica. [SE]

Global temperatures remained relatively steady during the Oligocene and early Miocene, and the Antarctic glaciation waned during that time. At around 15 Ma, subduction-related volcanism between central and South America created the connection between North and South America, preventing water from flowing between the Pacific and Atlantic Oceans. This further restricted the transfer of heat from the tropics to the poles, leading to a rejuvenation of the Antarctic glaciation. The expansion of that ice sheet increased Earth’s reflectivity enough to promote a positive feedback loop of further cooling: more reflective glacial ice, more cooling, more ice, etc. By the Pliocene (

5 Ma) ice sheets had started to grow in North America and northern Europe (Figure 16.5). The most intense part of the current glaciation — and the coldest climate — has been during the past million years (the last one-third of the Pleistocene), but if we count Antarctic glaciation, it really extends from the Oligocene to the Holocene, and will likely continue into the future.

The Pleistocene has been characterized by significant temperature variations (through a range of almost 10°C) on time scales of 40,000 to 100,000 years, and corresponding expansion and contraction of ice sheets. These variations are attributed to subtle changes in Earth’s orbital parameters (Milankovitch cycles), which are explained in more detail in Chapter 21. Over the past million years, the glaciation cycles have been approximately 100,000 years this variability is visible in Figure 16.5.

Figure 16.5 Foram oxygen isotope record for the past 5 million years based on O isotope data from sea-floor sediments [Created by SE using from data at http://www.lorraine-lisiecki.com/stack.html, Lisiecki and Raymo, 2005]

Exercise 16.1 Pleistocene Glacials and Interglacials

This diagram shows the past 500,000 years of the same data set as that shown in Figure 16.5. The last five glacial periods are marked with snowflakes. The most recent one, which peaked at around 20 ka, is known as the Wisconsin Glaciation. Describe the nature of temperature change that followed each of these glacial periods.

The current interglacial (Holocene) is marked with an H. Point out the previous five interglacial periods.

At the height of the last glaciation (Wisconsin Glaciation), massive ice sheets covered almost all of Canada and much of the northern United States (Figure 16.6). The massive Laurentide Ice Sheet covered most of eastern Canada, as far west as the Rockies, and the smaller Cordilleran Ice Sheet covered most of the western region. At various other glacial peaks during the Pleistocene and Pliocene, the ice extent was similar to this, and in some cases, even more extensive. The combined Laurentide and Cordilleran Ice Sheets were comparable in volume to the current Antarctic Ice Sheet.

Figure 16.6 The extent of the Cordilleran and Laurentide Ice Sheets near the peak of the Wisconsin Glaciation, around 15 ka. [redrawn by SE based on a map at: https://www.ncdc.noaa.gov/paleo/glaciation.html]


Interactive discussion

Status: closed

In this article, Zhou and colleagues report changes in winter glacier velocity for seven regions of Asia between 2017-2018 and 1999-2000. Glacier surface velocities are derived from satellite image correlation. They use two sensors: Landsat-7 (L7) for the period 1999-2000 and Sentinel-2 (S2) for the period 2017-2018. Contrary to previous studies, they observe glacier slowdown in the Karakoram and acceleration in the eastern part of the Himalaya. They find that thinning glaciers have increasing winter velocities and stable/thickening glaciers have decreasing winter velocities. The latter result is very surprising and seems in contradiction with the theoretical framework of ice dynamics understandings.

While the text is well written and the figures are of high quality, the methods are not precise enough to allow a replication of the work. I also suspect major flaws in the data processing and analysis that leads to erroneous results and conclusions. In particular, the interpretation of changes in velocities in relationship with changes in ice thickness is not convincing at all (i.e. increased accumulation suggested in regions where glaciers are losing mass). Below I provide general comments about these major flaws and some technical remarks on the text.

1 - The velocity changes are not reliable and reproducible.

Analyzing glacier velocity changes from different sensors is very challenging and requires to check some basic requirements that are not full-filled here. Dehecq et al. (2019) showed that using different sensors can introduce large biases in velocity. Please find below a check-list of critical methodological points that need to be addressed:

  • Calculate velocity changes on exactly the same pixels. From the text (L76-77), I understand that the authors calculate a mean velocity for each glacier for the first period, and then a mean velocity for the second period. The velocity change is calculated as the difference between these two terms. This framework is not suitable to calculate velocity changes, because the glacier velocity is highly variable in space, and consequently the mean glacier velocity for each period must be calculated exactly on the same pixels. This is even more critical when two different sensors are used (L7 and S2 here). These sensors have different capabilities and S2 likely produces reliable velocity fields for much larger fraction of the glacier surface than L7.
  • Due to the non-gaussian distribution of residuals (i.e. the fact that the modulus of the velocity vector is always positive), you need to find a way to normalize the velocity changes. One good check of the efficiency of the chosen normalization is that the velocity change on stable terrain should be zero. Please provide the velocity changes on stable terrain to demonstrate the robustness of your processing. On this topic, I recommend to read thoroughly the supplementary of Dehecq et al. (2019), and in particular the figures S8 and S9.

2 - The interpretation of the observed changes is not convincing

This article is very short, which is in general good for scientific writing in my opinion. However, here I feel that I am missing important parts of the message, due to the text’s lack of details. For instance, the authors should give more context/background about the section 4.1 (“Relationship between glacier surface velocity and geometry”). Why are these relationships investigated? Why is the change in velocity expected to be related to the slope, area or other variables?

I am also missing a more precise interpretation within the climate context. The authors study winter glacier velocities over a very large region with contrasted climate settings. In the central Himalaya and Nyainqentanglha glaciers accumulate during the monsoon/spring period, while in Karakoram they accumulate during winter (Maussion et al., 2014). As a consequence, winter does not mean accumulation period everywhere and any interpretation related to the climate context must be much finer than the proposed analysis.

The authors attribute the difference between their results and Dehecq’s results by the fact they measure winter velocities, whereas Dehecq et al. measured summer velocities. The interpretation of the authors is not right here: Dehecq et al. measured annual velocities (fig. S2 of Dehecq et al.), with the average annual velocity field centered roughly in summer (fig. S3 of Dehecq et al.). Seasonal glacier velocity variability is poorly documented, especially for Asian glaciers (Armstrong et al., 2017 Usman & Furuya, 2018), but I doubt that the difference between the two studies originates from it. I suggest that the authors apply their workflow to summer image pairs to demonstrate that they can replicate Dehecq’s results first.

L137-149 are very difficult to follow and do not make much sense to me. From my understanding, the negative relationship found in this study (fig. 6) lead to the conclusion that the parameter m in eq. 1 should be negative? This is in strong contradiction with basic physics of ice dynamics (e.g., Cuffey & Paterson, 2010). If the authors think that “ice mass loss promotes glacier motion in winter”, they need to suggest a mechanism that could explain this. I don’t follow the logics of the rest of the section (L145-149), and I think that most of these statements are not correct (and not backed up by any literature reference).

Title: “Himalayas” is not the name of the studied region, which encompass Karakoram, Hindu Kush, Himalaya and Nyainqêntanglha

Structure: it is clearer to write three separate sections for the methods, results and discussion. For instance, L90-100 read like a results sections and not a discussion.

L10-11: this is not shown in the paper

L55: this is incorrect, see my comments above

L70-71: a better way to evaluate the velocity changes accuracy would be to look at stable terrain changes

Armstrong, W. H., Anderson, R. S., & Fahnestock, M. A. (2017). Spatial Patterns of Summer Speedup on South Central Alaska Glaciers. Geophysical Research Letters, 44(18), 9379–9388. https://doi.org/10.1002/2017GL074370

Cuffey, K. M., & Paterson, W. S. B. (2010). The physics of glaciers. Academic Press.

Dehecq, A., Gourmelen, N., Gardner, A. S., Brun, F., Goldberg, D., Nienow, P. W., et al. (2019). Twenty-first century glacier slowdown driven by mass loss in High Mountain Asia. Nature Geoscience, 12(1), 22–27. https://doi.org/10.1038/s41561-018-0271-9

Maussion, F., Scherer, D., Mölg, T., Collier, E., Curio, J., & Finkelnburg, R. (2014). Precipitation Seasonality and Variability over the Tibetan Plateau as Resolved by the High Asia Reanalysis. Journal of Climate, 27(5), 1910–1927. https://doi.org/10.1175/JCLI-D-13-00282.1

Usman, M., & Furuya, M. (2018). Interannual modulation of seasonal glacial velocity variations in the Eastern Karakoram detected by ALOS-1/2 data. Journal of Glaciology, 64(245), 465–476.


Thread: IRIS WEBINARS: Glacial Earthquakes and Cryoseismology as a Tool for Investigating Greenland Outlet Glaciers, 2/21 at 2 PM Eastern

Please register for Glacial Earthquakes and Cryoseismology as a Tool for Investigating Greenland Outlet Glaciers on February 21, 2018 2:00 PM EST at:

Presenter: Dr. Stephen A. Veitch, University of Texas at El Paso

Abstract: Although seismology has been used as a tool to investigate glaciers and icesheets for decades, the field of 'Cryoseismology' has exploded in the last decade with the development of new tools and techniques that have helped to facilitate the growth of the field. In this talk we will cover ongoing and published research from Greenland involving both locally and globally recorded seismograms. The loss of ice from the Greenland ice sheet is an important contributor to current and future sea level rise occurring due to ongoing changes in the global climate. A significant portion of this ice mass loss comes through the calving of large icebergs at Greenland's many marine-terminating outlet glaciers. However, the dynamics of calving at these glaciers is currently not well understood, complicating projections of future behaviour of these glaciers and mass loss from the Greenland ice sheet.

On the global scale, data from the long-standing global seismic network has recorded the occurrence of glacial earthquakes, large long period earthquakes that occur during large calving events at near-grounded outlet glaciers. The occurrence and source parameters of these earthquakes provide insight into the link between glacier calving and climatic and oceanic forcings, as well as information on the large-scale glacier-dynamic conditions under which these major calving events occur. On the more local scale, a deployment of seismometers around an individual glacier has provided insights on the seismic environment of a calving glacier, as well as the more immediate, short-term external drivers of calving events. We consider both local and global seismic data in order to further understanding of the dynamics of the calving process at Greenland outlet glaciers, and find that glacial earthquake production is indicative of a near-grounded terminus at the source glacier. We find that the locations derived from these events are accurate and are sensitive to changes in the calving-front position of the source glacier, and that the active-force azimuths are representative of the orientation of the glacier at the time of calving. We also find that these glaciers are the source of abundant small icequakes, which are strongly tied to the occurrence of major calving events. The small icequakes that occur at Helheim glacier are modulated by semi-diurnal variations in tide height, and potentially control the timing of major calving events by progressively damaging the glacier tongue.

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21.2: Glaciers - Geosciences

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Glaciers, Sea-Ice and Permafrost

Our research group studies a wide range of cryospheric processes, with a particular emphasis on the relation between climate and changes in the cryosphere. We study questions from the modern contributions of glaciers and ice sheets to sea level rise to inferring past climate from ice core records. We study the changing Arctic sea ice from the micro-scale structure of the ice crystals to the basin-scale dynamic movement of the ice. We study the growth and decay of permafrost and its affect on environmental and engineering problems. We develop methods to physically measure and mathematically model the cryosphere to understand the physical processes and interactions with the climate system, we observe and monitor changes in the cryosphere, and we predict impacts of cryospheric change on the local and global environment (e.g. ecosystems, hydrology, carbon cycle, sea level, ocean dynamics) as well as human-related concerns (e.g. infrastructure). We are interested in all aspects of how the cryosphere is affected by global change with respect to climate as well as natural and human induced disturbances.

Faculty in this research are also affiliated with the Geophysical Institute's Snow, Ice, and Permafrost research group. More information on different aspects of cryospheric research can be found throught he links below:


Author information

Affiliations

Earth Science and Observation Center, Cooperative Institute for Research in Environmental Sciences, University of Colorado Boulder, Boulder, CO, USA

Department of Physics, Emory University, Atlanta, GA, USA

Department of Natural Sciences, University of Alaska Southeast, Juneau, AK, USA

Geophysical Institute, University of Alaska Fairbanks, Fairbanks, AK, USA

Mark A. Fahnestock & Martin Truffer

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Contributions

R.K.C., J.M.A., M.A.F. and M.T. collected the TRI data. R.K.C. processed and analysed the data with input from all collaborators. J.C.B. created and completed the modelling component. R.K.C., J.C.B. and J.M.A. authored the manuscript with input from M.A.F. and M.T.

Corresponding author


Department of Geosciences

Geosciences is the study the Earth and its people through geography, geology and geophysics. This includes a broad range of subjects like volcanism, plate tectonics, glaciology, remote sensing and mineralogy. Geography can focus on physical science or it can investigate the patterns and processes that shape human society like religion or economics. Geoscience jobs are plentiful, salaries are competitive, and the demand for young and enthusiastic geoscientists is expected to grow.

WHY STUDY GEOSCIENCE AT UAF AND IN ALASKA?

  • For students interested in geosciences, Alaska is one of the most exciting natural laboratories on Earth.
  • Alaska saw the world’s largest volcanic eruption and second largest earthquake during the last century.
  • Together with western Canada, Alaska contains the largest ice mass outside of Antarctica and Greenland.
  • Global change is warming Alaska and causing dramatic changes to permafrost, glaciers, and sea ice.
  • Alaska is an area of active exploration for and discovery of mineral and energy resources such as oil and rare earth minerals.
  • Our program emphasizes field experience with classes like the geology field camp in which students practice geological mapping in remote mountains.
  • Our research spans the globe with faculty work at research sites like Japan, Russia, Antarctica. One even studies the geology of Mars.

Our classes allow you to explore

  • Economic Geology
  • Earth Science
  • Geochronology (studying timescales embedded in rocks),
  • GIS
  • Glaciers
  • Lake, river & sea ice
  • Landscape analysis through geography
  • Paleontology
  • Permafrost
  • Natural hazards and their mitigation
  • Remote sensing
  • Rocks and minerals
  • Stratigraphy & sedimentation
  • Teaching geology at a middle or high school
  • Tectonics
  • Volcanoes
  • and more…

Career opportunities with a degree in the Geosciences

Some of the questions related to geosciences that UAF researchers are asking

Geography

What happened to the ice-age horse that inhabited Beringia, the biogeographic connector between Asia and North America

Geophysics

What’s happening to the Pacific Plate under the Wrangell Mountains? It’s as if the plate just disappears.


Watch the video: How do glaciers shape the landscape? Animation from Kerboodle.