[[Image:Antarctica glacier flow rate.jpg|300 px|thumb|Glacial flow rate in the Antarctic ice sheet.]]
[[Image:Antarctica glacier flow rate.jpg|300 px|thumb|Glacial flow rate in the Antarctic ice sheet.]]
[[File:Flow of Ice Across Antarctica.ogv|thumb|upright=1.35|The motion of ice in Antarctica]]
[[File:Flow of Ice Across Antarctica.ogv|thumb|upright=1.35|The motion of ice in Antarctica]]
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On an hour-to-hour basis, surges of ice motion can be modulated by tidal activity. The influence of a 1 m tidal oscillation can be felt as much as 100 km from the sea.<ref name=Clarke2005/> During larger [[spring tide]]s, 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.<ref name=Bindschalder2003>{{cite journal |last1=Bindschadler |first1=Robert A. |last2=King |first2=Matt A. |last3=Alley |first3=Richard B. |last4=Anandakrishnan |first4=Sridhar |last5=Padman |first5=Laurence |title=Tidally Controlled Stick-Slip Discharge of a West Antarctic Ice |journal=Science |date=22 August 2003 |volume=301 |issue=5636 |pages=1087–1089 |doi=10.1126/science.1087231 |pmid=12934005 |s2cid=37375591 |url=https://zenodo.org/record/1230832 }}</ref><ref name=Anandakrishnan2003>{{cite journal |last1=Anandakrishnan |first1=S. |last2=Voigt |first2=D. E. |last3=Alley |first3=R. B. |last4=King |first4=M. A. |title=Ice stream D flow speed is strongly modulated by the tide beneath the Ross Ice Shelf |journal=Geophysical Research Letters |date=April 2003 |volume=30 |issue=7 |page=1361 |doi=10.1029/2002GL016329 |bibcode=2003GeoRL..30.1361A |s2cid=53347069 |doi-access=free }}</ref> At neap tides, this interaction is less pronounced, and surges instead occur approximately every 12 hours.<ref name=Bindschalder2003/>
On an hour-to-hour basis, surges of ice motion can be modulated by tidal activity. The influence of a 1 m tidal oscillation can be felt as much as 100 km from the sea.<ref name=Clarke2005/> During larger [[spring tide]]s, 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.<ref name=Bindschalder2003>{{cite journal |last1=Bindschadler |first1=Robert A. |last2=King |first2=Matt A. |last3=Alley |first3=Richard B. |last4=Anandakrishnan |first4=Sridhar |last5=Padman |first5=Laurence |title=Tidally Controlled Stick-Slip Discharge of a West Antarctic Ice |journal=Science |date=22 August 2003 |volume=301 |issue=5636 |pages=1087–1089 |doi=10.1126/science.1087231 |pmid=12934005 |s2cid=37375591 |url=https://zenodo.org/record/1230832 }}</ref><ref name=Anandakrishnan2003>{{cite journal |last1=Anandakrishnan |first1=S. |last2=Voigt |first2=D. E. |last3=Alley |first3=R. B. |last4=King |first4=M. A. |title=Ice stream D flow speed is strongly modulated by the tide beneath the Ross Ice Shelf |journal=Geophysical Research Letters |date=April 2003 |volume=30 |issue=7 |page=1361 |doi=10.1029/2002GL016329 |bibcode=2003GeoRL..30.1361A |s2cid=53347069 |doi-access=free }}</ref> At neap tides, this interaction is less pronounced, and surges instead occur approximately every 12 hours.<ref name=Bindschalder2003/>
<!--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. -->
<!--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===
===Subglacial processes===
[[File:Glacier cross-section.jpg|thumb|upright|A cross-section through a glacier. The base of the glacier is more transparent as a result of melting.]]
[[File:Glacier cross-section.jpg|thumb|upright|A cross-section through a glacier. The base of the glacier is more transparent as a result of melting.]]
Revisionasof11:42,6April2024
Large mass of glacial ice
"Continental glacier" redirects here. For the glacier located in Wyoming, see Continental Glacier.
Inglaciology, an ice sheet, also known as a continental glacier,[2] is a mass of glacialice 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.
"Ice flow" redirects here. For floating ice, see Ice floe.
An ice sheet is a body of ice which has formed over thousands of years of snow accumulation and covers land area of continental size - i.e. >50,000 km2. Even stable ice sheets are continually in motion as the ice gradually flows outward from the central plateau, which is the tallest point of the ice sheet, and towards the margins. The ice sheet slope is low around the plateau but increases steeply at the margins.[4] This difference in slope occurs due to an imbalance between high ice accumulation in the central plateau and lower accumulation, as well as higher ablation, at the margins. 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.[5][6] This motion 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 (i.e. tidal flows) to the centennial (Milankovich cycles).[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:[5][6]
where:
= shear strain (flow) rate
= stress
= a constant between 2–4 (typically 3 for most glaciers) that increases with lower temperature
= a temperature-dependent constant
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.[5][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.[7]
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.[8] 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.[8]
On an hour-to-hour basis, surges of ice motion can be modulated by tidal activity. The influence of a 1 m tidal oscillation can be felt as much as 100 km from the sea.[9] 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, and surges instead occur approximately every 12 hours.[10]
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]
τD = ρgh sin α
where τD is the driving stress, and α the ice surface slope in radians.[9]
τB is the basal shear stress, a function of bed temperature and softness.[9]
τF, the shear stress, is the lower of τB and τD. It controls the rate of plastic flow, as per the figure (inset, right).
For a given glacier, the two variables are τD, which varies with h, the depth of the glacier, and τB, 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 τB).[9] If the sediment strength falls far below τD, movement of the glacier will be accommodated by motion in the sediments, as opposed to sliding.
Porosity may vary through a range of methods.
Movement of the overlying glacier may cause the bed to undergo dilatancy; the resulting shape change reorganises blocks. This reorganises closely packed blocks (a little like neatly folded, tightly packed clothes in a suitcase) into a messy jumble (just as clothes never fit back in when thrown in in a disordered fashion). This increases the porosity. Unless water is added, this will necessarily reduce the pore pressure (as the pore fluids have more space to occupy).[9]
Pressure may cause compaction and consolidation of underlying sediments.[9] Since water is relatively incompressible, this is easier when the pore space is filled with vapour; any water must be removed to permit compression. In soils, this is an irreversible process.[9]
Sediment degradation by abrasion and fracture decreases the size of particles, which tends to decrease pore space. However, the motion of the particles may disorder the sediment, with the opposite effect. These processes also generate heat.[9]
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.[12]
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.[12] Further, bed roughness can also 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]
As well as affecting the sediment stress, fluid pressure (pw) 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, pi, given by ρgh. Under fast-flowing ice streams, these two pressures will be approximately equal, with an effective pressure (pi – pw) 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 melting
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.[13] 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.[13] The subglacial motion of water is expressed in the surface topography of ice sheets, which slump down into vacated subglacial lakes.[13]
The presence of basal meltwater depends on both bed temperature and other factors. For instance, 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.[12]
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, τD 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.[12]
Increasing global air temperatures due to climate change take around 10,000 years to directly propagate through the ice before they influence bed temperatures, but may have an effect through increased surface melting, producing more supraglacial lakes. These lakes may feed warm water to glacial bases and facilitate glacial motion.[14] 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.[15] Water that reaches the bed of a glacier may freeze there, increasing the thickness of the glacier by pushing it up from below.[16]
Boundary conditions
As the margins end at the marine boundary, excess ice is discharged through ice streams or outlet glaciers. Then, it either falls directly into the sea or is accumulated atop the floating ice shelves.[4]: 2234 Those ice shelves then calve icebergs at their periphery if they experience excess of ice. Ice shelves would also experience accelerated calving due 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.[17]
The presence of ice shelves has a stabilizing influence on the glacier behind them, while an absence of an ice shelf becomes destabilizing. For instance, when Larsen B ice shelf in the Antarctic Peninsula had collapsed over three weeks in February 2002, the four glaciers behind it - Crane Glacier, Green Glacier, Hektoria Glacier and Jorum Glacier - all started to flow at a much faster rate, while the two glaciers (Flask and Leppard) stabilized by the remnants of the ice shelf did not accelerate.[18] The collapse of the Larsen B shelf was preceded by thinning of just 1 metre per year, while some other Antarctic ice shelves have displayed thinning of tens of metres per year.[14] Further, increased ocean temperatures of 1 °C may lead to up to 10 metres per year of basal melting.[14] 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.[14]
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.[19] 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.[19][20]
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.[21] 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.[21] 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.[20] 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.[20] 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.[20][22]
Vulnerable locations
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.[24][25] In particular, the Thwaites and Pine Island glaciers are most likely to be prone to MISI, and both glaciers have been rapidly thinning and accelerating in recent decades.[26][27][28][29] As the result, sea level rise from the ice sheet could be accelerated by tens of centimeters within the 21st century alone.[30]
The majority of the East Antarctic Ice Sheet would not be affected. Totten Glacier is the largest glacier there which is known to be subject to MISI - yet, its potential contribution to sea level rise is comparable to that of the entire West Antarctic Ice Sheet.[31] Totten Glacier has been losing mass nearly monotonically in recent decades,[32] 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.[33][34][35] The Wilkes Basin is the only major submarine basin in Antarctica that is not thought to be sensitive to warming.[28] Ultimately, even geologically rapid sea level rise would still most likely require several millennia for the entirety of these ice masses (WAIS and the subglacial basins) to be lost.[36][37]
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.[20] 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.[21][38] 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.[39]
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,[43][44] and a substantial retreat of its coastal glaciers since at least the 1990s.[45] Estimates suggest it added around 7.6 ± 3.9 mm (19⁄64 ± 5⁄32in) to the global sea level rise between 1992 and 2017,[46] and has been losing ice in the 2010s at a rate equivalent to 0.4 millimetres (0.016 inches) of annual sea level rise.[47] 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.[48] Fresh meltwater from WAIS also contributes to ocean stratification and dilutes the formation of salty Antarctic bottom water, which destabilizes Southern Ocean overturning circulation.[48][49][50]
In the long term, the West Antarctic Ice Sheet is likely to disappear due to the warming which has already occurred.[51]Paleoclimate evidence suggests that this has already happened during the Eemian period, when the global temperatures were similar to the early 21st century.[52][53] It is believed that the loss of the ice sheet would take place between 2,000 and 13,000 years in the future,[54][55] although several centuries of high emissions may shorten this to 500 years.[56] 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,[57] but this would require a higher level of warming.[58]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.[56]
The preservation of WAIS may require a persistent reduction of global temperatures to 1 °C (1.8 °F) below the preindustrial level, or to 2 °C (3.6 °F) below the temperature of 2020.[59] Because the collapse of the ice sheet would be preceded by the loss of Thwaites Glacier and Pine Island Glacier, some have instead proposed interventions to preserve them. In theory, adding thousands of gigatonnes of artificially created snow could stabilize them,[60] but it would be extraordinarily difficult and may not account for the ongoing acceleration of ocean warming in the area.[51] Others suggest that building obstacles to warm water flows beneath glaciers would be able to delay the disappearance of the ice sheet by many centuries, but it would still require one of the largest civil engineering interventions in history.
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[65] results in the reported cold temperature records of nearly −100 °C (−148 °F).[66][67] 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.[65]
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.[68][69] 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.[70][71][72] 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.[73]
Because the East Antarctic ice sheet has barely warmed, it is still gaining ice on average.[74][75] for instance, GRACE satellite data indicated East Antarctica mass gain of 60 ± 13 billion tons per year between 2002 and 2010.[76] It is most likely to first see sustained losses of ice at its most vulnerable locations such as Totten Glacier and Wilkes Basin. Those areas are sometimes collectively described as East Antarctica's subglacial basins, and it is believed that once the warming reaches around 3 °C (5.4 °F), then they would start to collapse over a period of around 2,000 years,[77][78] This collapse would ultimately add between 1.4 m (4 ft 7 in) and 6.4 m (21 ft 0 in) to sea levels, depending on the ice sheet model used.[79] The EAIS as a whole holds enough ice to raise global sea levels by 53.3 m (175 ft).[64] However, it would take global warming in a range between 5 °C (9.0 °F) and 10 °C (18 °F), and a minimum of 10,000 years for the entire ice sheet to be lost.[77][78]
The Greenland ice sheet is an ice sheet which forms the second largest body of ice in the world. It is an average of 1.67 km (1.0 mi) thick, and over 3 km (1.9 mi) thick at its maximum.[80] It is almost 2,900 kilometres (1,800 mi) long in a north–south direction, with a maximum width of 1,100 kilometres (680 mi) at a latitude of 77°N, near its northern edge.[81] The ice sheet covers 1,710,000 square kilometres (660,000 sq mi), around 80% of the surface of Greenland, or about 12% of the area of the Antarctic ice sheet.[80] The term 'Greenland ice sheet' is often shortened to GIS or GrIS in scientific literature.[82][83][84][85]
Greenland has had major glaciers and ice caps for at least 18 million years,[86] but a single ice sheet first covered most of the island some 2.6 million years ago.[87] Since then, it has both grown[88][89] and contracted significantly.[90][91][92] The oldest known ice on Greenland is about 1 million years old.[93] Due to anthropogenic greenhouse gas emissions, the ice sheet is now the warmest it has been in the past 1000 years,[94] and is losing ice at the fastest rate in at least the past 12,000 years.[95]
Every summer, parts of the surface melt and ice cliffs calve into the sea. Normally the ice sheet would be replenished by winter snowfall,[83] but due to global warming the ice sheet is melting two to five times faster than before 1850,[96] and snowfall has not kept up since 1996.[97] 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 (2+1⁄2in) 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,[98]: 1302 with a worst-case of about 33 cm (13 in).[99] For comparison, melting has so far contributed 1.4 cm (1⁄2in) since 1972,[100] while sea level rise from all sources was 15–25 cm (6–10 in)) between 1901 and 2018.[101]: 5
If all 2,900,000 cubic kilometres (696,000 cu mi) of the ice sheet were to melt, it would increase global sea levels by ~7.4 m (24 ft).[80] Global warming between 1.7 °C (3.1 °F) and 2.3 °C (4.1 °F) would likely make this melting inevitable.[85] However, 1.5 °C (2.7 °F) would still cause ice loss equivalent to 1.4 m (4+1⁄2ft) of sea level rise,[102] and more ice will be lost if the temperatures exceed that level before declining.[85] If global temperatures continue to rise, the ice sheet will likely disappear within 10,000 years.[103][104] At very high warming, its future lifetime goes down to around 1,000 years.[99]
Carbon cycle
Historically, ice sheets were viewed as inert components of the carbon cycle and were largely disregarded in global models. In 2010s, research had demonstrated 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.[105] There is a massive hemispheric contrast in carbon storage between the two ice sheets. While only about 0.5-27 billion tonnes of pure carbon are present underneath the Greenland ice sheet, 6000-21,000 billion tonnes are thought to be located underneath Antarctica.[105] For comparison, 1400–1650 billion tonnes are contained within the Arctic permafrost,[106] while the annual anthropogenic emissions amount to around 40 billion tonnes of CO2.[30]: 1237 ) This carbon can act as a climate change feedback if it is gradually released through meltwater, thus increasing overall carbon dioxide emissions.[107]
In Greenland, there is one known area, at Russell Glacier, where meltwater carbon is released into the atmosphere as methane, which has a much larger global warming potential than carbon dioxide:[108] however, it also harbours large numbers of methanotrophic bacteria, which limit those emissions.[109][110]
In geologic timescales
Normally, the transitions between glacial and interglacial states are governed by Milankovitch cycles, which are patterns in insolation (the amount of sunlight reaching the Earth). These patterns are caused by the variations in shape of the Earth's orbit and its angle relative to the Sun, caused by the gravitational pull of other planets as they go through their own orbits.[111][112]
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.[113]
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.[114]
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.[115][116]
The icing of Antarctica began in the Late Palaeocene or middle Eocene between 60[117] and 45.5 million years ago[118] and escalated during the Eocene–Oligocene extinction event about 34 million years ago. CO2 levels were then about 760 ppm[119] and had been decreasing from earlier levels in the thousands of ppm. Carbon dioxide decrease, with a tipping point of 600 ppm, was the primary agent forcing Antarctic glaciation.[120] The glaciation was favored by an interval when the Earth's orbit favored cool summers but oxygen isotope ratio cycle marker changes were too large to be explained by Antarctic ice-sheet growth alone indicating an ice age of some size.[121] The opening of the Drake Passage may have played a role as well[122] though models of the changes suggest declining CO2 levels to have been more important.[123]
The Western Antarctic ice sheet declined somewhat during the warm early Pliocene epoch, approximately five to three million years ago; during this time the Ross Sea opened up.[124] But there was no significant decline in the land-based Eastern Antarctic ice sheet.[125]
While there is evidence of large glaciersinGreenland for most of the past 18 million years,[86] these ice bodies were probably similar to various smaller modern examples, such as Maniitsoq and Flade Isblink, which cover 76,000 and 100,000 square kilometres (29,000 and 39,000 sq mi) around the periphery. Conditions in Greenland were not initially suitable for a single coherent ice sheet to develop, but this began to change around 10 million years ago, during the middle Miocene, when the two passive continental margins which now form the uplands of West and East Greenland experienced uplift, and ultimately formed the upper planation surface at a height of 2000 to 3000 meter above sea level.[126][127]
Later uplift, during the Pliocene, formed a lower planation surface at 500 to 1000 meters above sea level. A third stage of uplift created multiple valleys and fjords below the planation surfaces. This uplift intensified glaciation due to increased orographic precipitation and cooler surface temperatures, allowing ice to accumulate and persist.[126][127] As recently as 3 million years ago, during the Pliocene warm period, Greenland's ice was limited to the highest peaks in the east and the south.[128] Ice cover gradually expanded since then,[87] until the atmospheric CO2 levels dropped to between 280 and 320 ppm 2.7–2.6 million years ago, by which time temperatures had dropped sufficiently for the disparate ice caps to connect and cover most of the island.[82]
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