Ice sheet: Difference between revisions

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Inglaciology, an ice sheet, also known as a continental glacier,^[2] is a mass of glacial ice that covers surrounding terrain and is greater than 50,000 km² (19,000 sq mi).^[3] The only current ice sheets are the Antarctic ice sheet and the Greenland ice sheet. Ice sheets are bigger than ice shelves or alpine glaciers. Masses of ice covering less than 50,000 km² are termed an ice cap. An ice cap will typically feed a series of glaciers around its periphery.

Although the surface is cold, the base of an ice sheet is generally warmer due to geothermal heat. In places, melting occurs and the melt-water lubricates the ice sheet so that it flows more rapidly. This process produces fast-flowing channels in the ice sheet — these are ice streams.

Definition

An ice sheet is "an ice body originating on land that covers an area of continental size, generally defined as covering >50,000 km² , and that has formed over thousands of years through accumulation and compaction of snow".^[4]^: 2234

Common properties

Ice sheets have the following properties: "An ice sheet flows outward from a high central ice plateau with a small average surface slope. The margins usually slope more steeply, and most ice is discharged through fast-flowing ice streams or outlet glaciers, often into the sea or into ice shelves floating on the sea."^[4]^: 2234

Ice movement is dominated by the motion of glaciers, whose activity is determined by a number of processes.^[6] Their motion is the result of cyclic surges interspersed with longer periods of inactivity, on both hourly and centennial time scales.

Until recently, ice sheets were viewed as inert components of the carbon cycle and were largely disregarded in global models. Research in the past decade has transformed this view, demonstrating the existence of uniquely adapted microbial communities, high rates of biogeochemical/physical weathering in ice sheets and storage and cycling of organic carbon in excess of 100 billion tonnes, as well as nutrients (see diagram).^[5]

Dynamics

The motion of ice sheets is driven by gravity but is controlled by temperature and the strength of individual glacier bases. A number of processes alter these two factors, resulting in cyclic surges of activity interspersed with longer periods of inactivity, on time scales ranging from hourly to the centennial.

Boundary conditions

The interface between an ice stream and the ocean is a significant control of the rate of flow.

Ice shelves are thick layers of ice floating on the sea – can stabilise the glaciers that feed them. These tend to have accumulation on their tops, may experience melting on their bases, and calve icebergs at their periphery. The catastrophic collapse of the Larsen B ice shelf in the space of three weeks during February 2002 yielded some unexpected observations. The glaciers that had fed the ice sheet (Crane, Jorum, Green, Hektoria – see image) increased substantially in velocity. This cannot have been due to seasonal variability, as glaciers flowing into the remnants of the ice shelf (Flask, Leppard) did not accelerate.^[7]

Ice shelves exert a dominant control in Antarctica, but are less important in Greenland, where the ice sheet meets the sea in fjords. Here, melting is the dominant ice removal process,^[8] resulting in predominant mass loss occurring towards the edges of the ice sheet, where icebergs are calved in the fjords and surface meltwater runs into the ocean.

Tidal effects are also important; the influence of a 1 m tidal oscillation can be felt as much as 100 km from the sea.^[9] On an hour-to-hour basis, surges of ice motion can be modulated by tidal activity. During larger spring tides, an ice stream will remain almost stationary for hours at a time, before a surge of around a foot in under an hour, just after the peak high tide; a stationary period then takes hold until another surge towards the middle or end of the falling tide.^[10]^[11] At neap tides, this interaction is less pronounced, without tides surges would occur more randomly, approximately every 12 hours.^[10]

Ice shelves are also sensitive to basal melting. In Antarctica, this is driven by heat fed to the shelf by the circumpolar deep water current, which is 3 °C above the ice's melting point.^[12]

As well as heat, the sea can also exchange salt with the oceans. The effect of latent heat, resulting from melting of ice or freezing of sea water, also has a role to play. The effects of these, and variability in snowfall and base sea level combined, account for around 80 mm annual variability in ice shelf thickness.

Long-term changes

Over long time scales, ice sheet mass balance is governed by the amount of sunlight reaching the Earth. This variation in sunlight reaching the Earth, or insolation, over geologic time is in turn determined by the angle of the Earth to the Sun and shape of the Earth's orbit, as it is pulled on by neighboring planets; these variations occur in predictable patterns called Milankovitch cycles. Milankovitch cycles dominate climate on the glacial–interglacial timescale, but there exist variations in ice sheet extent that are not linked directly with insolation.

For instance, during at least the last 100,000 years, portions of the ice sheet covering much of North America, the Laurentide Ice Sheet broke apart sending large flotillas of icebergs into the North Atlantic. When these icebergs melted they dropped the boulders and other continental rocks they carried, leaving layers known as ice rafted debris. These so-called Heinrich events, named after their discoverer Hartmut Heinrich, appear to have a 7,000–10,000-year periodicity, and occur during cold periods within the last interglacial.^[13]

Internal ice sheet "binge-purge" cycles may be responsible for the observed effects, where the ice builds to unstable levels, then a portion of the ice sheet collapses. External factors might also play a role in forcing ice sheets. Dansgaard–Oeschger events are abrupt warmings of the northern hemisphere occurring over the space of perhaps 40 years. While these D–O events occur directly after each Heinrich event, they also occur more frequently – around every 1500 years; from this evidence, paleoclimatologists surmise that the same forcings may drive both Heinrich and D–O events.^[14]

Hemispheric asynchrony in ice sheet behavior has been observed by linking short-term spikes of methane in Greenland ice cores and Antarctic ice cores. During Dansgaard–Oeschger events, the northern hemisphere warmed considerably, dramatically increasing the release of methane from wetlands, that were otherwise tundra during glacial times. This methane quickly distributes evenly across the globe, becoming incorporated in Antarctic and Greenland ice. With this tie, paleoclimatologists have been able to say that the ice sheets on Greenland only began to warm after the Antarctic ice sheet had been warming for several thousand years. Why this pattern occurs is still open for debate.^[15]^[16]

Glacier flows

The main cause of flow within glaciers can be attributed to an increase in the surface slope, brought upon by an imbalance between the amounts of accumulation vs. ablation. This imbalance increases the shear stress on a glacier until it begins to flow. The flow velocity and deformation will increase as the equilibrium line between these two processes is approached, but are also affected by the slope of the ice, the ice thickness and temperature.^[17]^[6]

When the amount of strain (deformation) is proportional to the stress being applied, ice will act as an elastic solid. Ice will not flow until it has reached a thickness of 30 meters (98 ft), but after 50 meters (164 ft), small amounts of stress can result in a large amount of strain, causing the deformation to become a plastic flow rather than elastic. At this point the glacier will begin to deform under its own weight and flow across the landscape. According to the Glen–Nye flow law, the relationship between stress and strain, and thus the rate of internal flow, can be modeled as follows:^[17]^[6]

The lowest velocities are near the base of the glacier and along valley sides where friction acts against flow, causing the most deformation. Velocity increases inward toward the center line and upward, as the amount of deformation decreases. The highest flow velocities are found at the surface, representing the sum of the velocities of all the layers below.^[17]^[6]

Glaciers may also move by basal sliding, where the base of the glacier is lubricated by meltwater, allowing the glacier to slide over the terrain on which it sits. Meltwater may be produced by pressure-induced melting, friction or geothermal heat. The more variable the amount of melting at surface of the glacier, the faster the ice will flow.^[18]

The top 50 meters of the glacier form the fracture zone, where ice moves as a single unit. Cracks form as the glacier moves over irregular terrain, which may penetrate the full depth of the fracture zone.

Subglacial processes

Most of the important processes controlling glacial motion occur in the ice-bed contact—even though it is only a few meters thick.^[9] Glaciers will move by sliding when the basal shear stress drops below the shear resulting from the glacier's weight.^{[clarification needed]}

For a given glacier, the two variables are τ_D, which varies with h, the depth of the glacier, and τ_B, the basal shear stress.^{[clarification needed]}

Basal shear stress

The basal shear stress is a function of three factors: the bed's temperature, roughness and softness.^[9]

Whether a bed is hard or soft depends on the porosity and pore pressure; higher porosity decreases the sediment strength (thus increases the shear stress τ_B).^[9] If the sediment strength falls far below τ_D, movement of the glacier will be accommodated by motion in the sediments, as opposed to sliding. Porosity may vary through a range of methods.

A soft bed, with high porosity and low pore fluid pressure, allows the glacier to move by sediment sliding: the base of the glacier may even remain frozen to the bed, where the underlying sediment slips underneath it like a tube of toothpaste. A hard bed cannot deform in this way; therefore the only way for hard-based glaciers to move is by basal sliding, where meltwater forms between the ice and the bed itself.^[19]

Bed softness may vary in space or time, and changes dramatically from glacier to glacier. An important factor is the underlying geology; glacial speeds tend to differ more when they change bedrock than when the gradient changes.^[19]

As well as affecting the sediment stress, fluid pressure (p_w) can affect the friction between the glacier and the bed. High fluid pressure provides a buoyancy force upwards on the glacier, reducing the friction at its base. The fluid pressure is compared to the ice overburden pressure, p_i, given by ρgh. Under fast-flowing ice streams, these two pressures will be approximately equal, with an effective pressure (p_i – p_w) of 30 kPa; i.e. all of the weight of the ice is supported by the underlying water, and the glacier is afloat.^[9]

Basal melt

A number of factors can affect bed temperature, which is intimately associated with basal meltwater. The melting point of water decreases under pressure, meaning that water melts at a lower temperature under thicker glaciers.^[9] This acts as a "double whammy", because thicker glaciers have a lower heat conductance, meaning that the basal temperature is also likely to be higher.^[19]

Bed temperature tends to vary in a cyclic fashion. A cool bed has a high strength, reducing the speed of the glacier. This increases the rate of accumulation, since newly fallen snow is not transported away. Consequently, the glacier thickens, with three consequences: firstly, the bed is better insulated, allowing greater retention of geothermal heat. Secondly, the increased pressure can facilitate melting. Most importantly, τ_D is increased. These factors will combine to accelerate the glacier. As friction increases with the square of velocity, faster motion will greatly increase frictional heating, with ensuing melting – which causes a positive feedback, increasing ice speed to a faster flow rate still: west Antarctic glaciers are known to reach velocities of up to a kilometre per year.^[9] Eventually, the ice will be surging fast enough that it begins to thin, as accumulation cannot keep up with the transport. This thinning will increase the conductive heat loss, slowing the glacier and causing freezing. This freezing will slow the glacier further, often until it is stationary, whence the cycle can begin again.^[19]

Supraglacial lakes represent another possible supply of liquid water to the base of glaciers, so they can play an important role in accelerating glacial motion. Lakes of a diameter greater than ~300 m are capable of creating a fluid-filled crevasse to the glacier/bed interface. When these crevasses form, the entirety of the lake's (relatively warm) contents can reach the base of the glacier in as little as 2–18 hours – lubricating the bed and causing the glacier to surge.^[20] Water that reaches the bed of a glacier may freeze there, increasing the thickness of the glacier by pushing it up from below.^[21]

Finally, bed roughness can act to slow glacial motion. The roughness of the bed is a measure of how many boulders and obstacles protrude into the overlying ice. Ice flows around these obstacles by melting under the high pressure on their stoss side; the resultant meltwater is then forced into the cavity arising in their lee side, where it re-freezes.^[9]

Pipe and sheet flow

The flow of water under the glacial surface can have a large effect on the motion of the glacier itself. Subglacial lakes contain significant amounts of water, which can move fast: cubic kilometres can be transported between lakes over the course of a couple of years.^[22]

This motion is thought to occur in two main modes: pipe flow involves liquid water moving through pipe-like conduits, like a sub-glacial river; sheet flow involves motion of water in a thin layer. A switch between the two flow conditions may be associated with surging behaviour. Indeed, the loss of sub-glacial water supply has been linked with the shut-down of ice movement in the Kamb ice stream.^[22] The subglacial motion of water is expressed in the surface topography of ice sheets, which slump down into vacated subglacial lakes.^[22]

Climate change

The implications of the current climate change on ice sheets are difficult to ascertain. It is clear that increasing temperatures are resulting in reduced ice volumes globally.^[8] (Due to increased precipitation, the mass of parts of the Antarctic ice sheet may currently be increasing, but the total mass balance is unclear.^[8])

Rising sea levels will reduce the stability of ice shelves, which have a key role in reducing glacial motion. Some Antarctic ice shelves are currently thinning by tens of metres per year, and the collapse of the Larsen B shelf was preceded by thinning of just 1 metre per year.^[8] Further, increased ocean temperatures of 1 °C may lead to up to 10 metres per year of basal melting.^[8] Ice shelves are always stable under mean annual temperatures of −9 °C, but never stable above −5 °C; this places regional warming of 1.5 °C, as preceded the collapse of Larsen B, in context.^[8]

Increasing global air temperatures take around 10,000 years to directly propagate through the ice before they influence bed temperatures, but may have an effect through increased surfacal melting, producing more supraglacial lakes, which may feed warm water to glacial bases and facilitate glacial motion.^[8] In areas of increased precipitation, such as Antarctica, the addition of mass will increase rate of glacial motion, hence the turnover in the ice sheet. Observations, while currently limited in scope, do agree with these predictions of an increasing rate of ice loss from both Greenland and Antarctica.^[8] A possible positive feedback may result from shrinking ice caps, in volcanically active Iceland at least. Isostatic rebound may lead to increased volcanic activity, causing basal warming – and, through CO₂ release, further climate change.^[23]

Cold meltwater provides cooling of the ocean's surface layer, acting like a lid, and also affecting deeper waters by increasing subsurface ocean warming and thus facilitating ice melt.

Erosion

Because ice can flow faster where it is thicker, the rate of glacier-induced erosion is directly proportional to the thickness of overlying ice. Consequently, pre-glacial low hollows will be deepened and pre-existing topography will be amplified by glacial action, while nunataks, which protrude above ice sheets, barely erode at all – erosion has been estimated as 5 m per 1.2 million years.^[25] This explains, for example, the deep profile of fjords, which can reach a kilometer in depth as ice is topographically steered into them. The extension of fjords inland increases the rate of ice sheet thinning since they are the principal conduits for draining ice sheets. It also makes the ice sheets more sensitive to changes in climate and the ocean.^[25]

Marine ice sheet instability

In the 1970s, Johannes Weertman proposed that because seawater is denser than ice, then any ice sheets grounded below sea level inherently become less stable as they melt due to Archimedes' principle.^[26] Effectively, these marine ice sheets must have enough mass to exceed the mass of the seawater displaced by the ice, which requires excess thickness. As the ice sheet melts and becomes thinner, the weight of the overlying ice decreases. At a certain point, sea water could force itself into the gaps which form at the base of the ice sheet, and marine ice sheet instability (MISI) would occur.^[26]^[27]

Even if the ice sheet is grounded below the sea level, MISI cannot occur as long as there is a stable ice shelf in front of it.^[28] The boundary between the ice sheet and the ice shelf, known as the grounding line, is particularly stable if it is constrained in an embayment.^[28] In that case, the ice sheet may not be thinning at all, as the amount of ice flowing over the grounding line would be likely to match the annual accumulation of ice from snow upstream.^[27] Otherwise, ocean warming at the base of an ice shelf tends to thin it through basal melting. As the ice shelf becomes thinner, it exerts less of an buttressing effect on the ice sheet, the so-called back stress increases and the grounding line is pushed backwards.^[27] The ice sheet is likely to start losing more ice from the new location of the grounding line and so become lighter and less capable of displacing seawater. This eventually pushes the grounding line back even further, creating a self-reinforcing mechanism.^[27]^[29]

Because the entire West Antarctic Ice Sheet is grounded below the sea level, it would be vulnerable to geologically rapid ice loss in this scenario.^[30]^[31] Sea level rise from the ice sheet could be accelerated by tens of centimeters within the 21st century alone.^[32] The majority of the East Antarctic Ice Sheet would not be affected, but its subglacial basins such as Wilkes Basin and the Aurora Subglacial Basin are also grounded below sea level and so subject to MISI. However, even geologically rapid sea level rise would still most likely require several millennia for the entirety of these ice masses to be lost.^[33]^[34]

Marine Ice Cliff Instability

A related process known as Marine Ice Cliff Instability (MICI) posits that due to the physical characteristics of ice, subaerial ice cliffs exceeding ~90 meters in height are likely to collapse under their own weight, and this could lead to runaway ice sheet retreat in a fashion similar to MISI.^[27] For an ice sheet grounded below sea level with an inland-sloping bed, ice cliff failure removes peripheral ice, which then exposes taller, more unstable ice cliffs, further perpetuating the cycle of ice front failure and retreat. Surface melt can further enhance MICI through ponding and hydrofracture.^[28]^[35] However, this process is considered more speculative than MISI, as it has never been observed at any scale. Some of the more detailed modelling has ruled it out.^[36]

Ocean warming

According to a 2016 published study, cold meltwater provides cooling of the ocean's surface layer, acting like a lid, and also affecting deeper waters by increasing subsurface ocean warming and thus facilitating ice melt.

Another theory discussed in 2007 for increasing warm bottom water is that changes in air circulation patterns have led to increased upwelling of warm, deep ocean water along the coast of Antarctica and that this warm water has increased melting of floating ice shelves.^[37] An ocean model has shown how changes in winds can help channel the water along deep troughs on the sea floor, toward the ice shelves of outlet glaciers.^[38]

Observations

In West Antarctica, the Thwaites and Pine Island glaciers have been identified to be potentially prone to MISI, and both glaciers have been rapidly thinning and accelerating in recent decades.^[39]^[40]^[41]^[42] In East Antarctica, Totten Glacier is the largest glacier known to be subject to MISI,^[43] and its potential contribution to sea level rise is comparable to that of the entire West Antarctic Ice Sheet. Totten Glacier has been losing mass nearly monotonically in recent decades,^[44] suggesting rapid retreat is possible in the near future, although the dynamic behavior of Totten Ice Shelf is known to vary on seasonal to interannual timescales.^[45]^[46]^[47] The Wilkes Basin is the only major submarine basin in Antarctica that is not thought to be sensitive to warming.^[41]

In October 2023, a study published in Nature Climate Change projected that ocean warming at about triple the historical rate is likely unavoidable in the 21st century, with no significant difference between mid-range emissions scenarios versus achieving the most ambitious targets of the Paris Agreement—suggesting that greenhouse gas mitigation has limited ability to prevent collapse of the West Antarctic Ice Sheet.^[48]

Earth's current two ice sheets

Antarctic ice sheet

West Antarctic ice sheet

West Antarctic ice sheet

Type	Ice sheet
Area	<1,970,000 km² (760,000 sq mi)^[49]
Thickness	~1.05 km (0.7 mi) (average),^[50] ~2 km (1.2 mi) (maximum)^[49]
Status	Receding

As a smaller part of Antarctica, WAIS is also more strongly affected by climate change. There has been warming over the ice sheet since the 1950s,^[52]^[53] and a substantial retreat of its coastal glaciers since at least the 1990s.^[54] Estimates suggest it added around 7.6 ± 3.9 mm (19⁄64 ± 5⁄32 in) to the global sea level rise between 1992 and 2017,^[55] and has been losing ice in the 2010s at a rate equivalent to 0.4 millimetres (0.016 inches) of annual sea level rise.^[56] While some of its losses are offset by the growth of the East Antarctic ice sheet, Antarctica as a whole will most likely lose enough ice by 2100 to add 11 cm (4.3 in) to sea levels. Further, marine ice sheet instability may increase this amount by tens of centimeters, particularly under high warming.^[57] Fresh meltwater from WAIS also contributes to ocean stratification and dilutes the formation of salty Antarctic bottom water, which destabilizes Southern Ocean overturning circulation.^[57]^[58]^[59]

In the long term, the West Antarctic Ice Sheet is likely to disappear due to the warming which has already occurred.^[60] Paleoclimate evidence suggests that this has already happened during the Eemian period, when the global temperatures were similar to the early 21st century.^[61]^[62] It is believed that the loss of the ice sheet would take place between 2,000 and 13,000 years in the future,^[63]^[64] although several centuries of high emissions may shorten this to 500 years.^[65] 3.3 m (10 ft 10 in) of sea level rise would occur if the ice sheet collapses but leaves ice caps on the mountains behind. Total sea level rise from West Antarctica increases to 4.3 m (14 ft 1 in) if they melt as well,^[66] but this would require a higher level of warming.^[67] Isostatic rebound of ice-free land may also add around 1 m (3 ft 3 in) to the global sea levels over another 1,000 years.^[65]

East Antarctic ice sheet

East Antarctic ice sheet

Type	Ice sheet
Thickness	~2.2 km (1.4 mi) (average),^[70] ~4.9 km (3.0 mi) (maximum) ^[71]

The surface of the EAIS is the driest, windiest, and coldest place on Earth. Lack of moisture in the air, high albedo from the snow as well as the surface's consistently high elevation^[74] results in the reported cold temperature records of nearly −100 °C (−148 °F).^[75]^[76] It is the only place on Earth cold enough for atmospheric temperature inversion to occur consistently. That is, while the atmosphere is typically warmest near the surface and becomes cooler at greater elevation, atmosphere during the Antarctic winter is cooler at the surface than in its middle layers. Consequently, greenhouse gases actually trap heat in the middle atmosphere and reduce its flow towards the surface while the temperature inversion lasts.^[74]

Due to these factors, East Antarctica had experienced slight cooling for decades while the rest of the world warmed as the result of climate change. Clear warming over East Antarctica only started to occur since the year 2000, and was not conclusively detected until the 2020s.^[77]^[78] In the early 2000s, cooling over East Antarctica seemingly outweighing warming over the rest of the continent was frequently misinterpreted by the media and occasionally used as an argument for climate change denial.^[79]^[80]^[81] After 2009, improvements in Antarctica's instrumental temperature record have proven that the warming over West Antarctica resulted in consistent net warming across the continent since the 1957.^[82]

Greenland ice sheet

Greenland has had major glaciers and ice caps for at least 18 million years,^[95] but a single ice sheet first covered most of the island some 2.6 million years ago.^[96] Since then, it has both grown^[97]^[98] and contracted significantly.^[99]^[100]^[101] The oldest known ice on Greenland is about 1 million years old.^[102] Due to anthropogenic greenhouse gas emissions, the ice sheet is now the warmest it has been in the past 1000 years,^[103] and is losing ice at the fastest rate in at least the past 12,000 years.^[104]

Every summer, parts of the surface melt and ice cliffs calve into the sea. Normally the ice sheet would be replenished by winter snowfall,^[92] but due to global warming the ice sheet is melting two to five times faster than before 1850,^[105] and snowfall has not kept up since 1996.^[106] If the Paris Agreement goal of staying below 2 °C (3.6 °F) is achieved, melting of Greenland ice alone would still add around 6 cm (21⁄2 in) to global sea level rise by the end of the century. If there are no reductions in emissions, melting would add around 13 cm (5 in) by 2100,^[107]^: 1302 with a worst-case of about 33 cm (13 in).^[108] For comparison, melting has so far contributed 1.4 cm (1⁄2 in) since 1972,^[109] while sea level rise from all sources was 15–25 cm (6–10 in)) between 1901 and 2018.^[110]^: 5

Melting due to climate change

Future melt of the West Antarctic ice sheet is potentially abrupt under a high emission scenario, as a consequence of a partial collapse.^[115]^{: 595–596} Part of the ice sheet is grounded on bedrock below sea level. This makes it possibly vulnerable to the self-enhancing process of marine ice sheet instability. Marine ice cliff instability could also contribute to a partial collapse. But there is limited evidence for its importance.^[114]^{: 1269–1270} A partial collapse of the ice sheet would lead to rapid sea level rise and a local decrease in ocean salinity. It would be irreversible for decades and possibly even millennia.^[115]^{: 595–596} The complete loss of the West Antarctic ice sheet would cause over 5 metres (16 ft) of sea level rise.^[116]

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