Editing IPCC:AR6/WGI/Chapter-9 (section)

=== 9.4.2 Antarctic Ice Sheet ===

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==== 9.4.2.1 Recent Observed Changes ====

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As stated in [[IPCC:Wg1:Chapter:Chapter-2#2.3.2.4|Section 2.3.2.4]] , satellite observations by Ice Sheet Mass Balance Intercomparison Exercise (IMBIE) combining multi-team estimates based on altimetry, gravity anomalies (GRACE) and the input-output method, already presented in SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ), are updated and extended to 2020 ( [[#The%20IMBIE%20Team--2021|The IMBIE Team, 2021]] ). The Antarctic Ice Sheet (AIS) lost 2670 [1800 to 3540] Gt mass over the period 1992–2020, equivalent to 7.4 [5.0 to 9.8] mm GMSL rise (for contribution to sea level budget, see Figures 9.16 and 9.18, and Table 9.5). Within uncertainties, this estimate agrees with a review of post-AR5 studies up to 2016 ( [[#Bamber--2018b|Bamber et al., 2018b]] ) and is consistent with recent single studies based on satellite laser altimetry ( [[#Smith--2020|Smith et al., 2020]] ), the input-output method ( [[#Rignot--2019|Rignot et al., 2019]] ) and gravimetry ( [[#Velicogna--2020|Velicogna et al., 2020]] ). The mass-loss rate was on average 49 [–2 to 100] Gt yr <sup>–1</sup> over the period 1992–1999, 70 [22 to 119] Gt yr <sup>–1</sup> over the period 2000–2009, and 148 [94 to 202] Gt yr <sup>–1</sup> over the period 2010–2016 (see Figures 9.16 and 9.18, and Table 9.SM.1). However, recent work suggests that the mass loss has not further increased since 2016 because of regional mass gains in Dronning Maud Land ( [[#Velicogna--2020|Velicogna et al., 2020]] ). Mass loss of the West Antarctic and Antarctic Peninsula ice sheets has increased since about 2000 ( ''very high confidence'' ), essentially due to increased ice discharge ( [[#Harig--2015|Harig and Simons, 2015]] ; [[#Paolo--2015|Paolo et al., 2015]] ; [[#Forsberg--2017|Forsberg et al., 2017]] ; [[#Bamber--2018b|Bamber et al., 2018b]] ; [[#Gardner--2018|Gardner et al., 2018]] ; [[#The%20IMBIE%20Team--2018|The IMBIE Team, 2018]] ; [[#Rignot--2019|Rignot et al., 2019]] ).

The SROCC reported with ''very high confidence'' that the acceleration, retreat and thinning of the principal West Antarctic outlet glaciers has dominated the observed Antarctic mass loss over the last decades, and stated with ''high confidence'' that these losses were driven by melting of ice shelves by warm ocean waters. The average West Antarctic Ice Sheet (WAIS) mass loss of 82 ± 9 Gt yr <sup>–1</sup> between 1992 and 2017 ( [[#The%20IMBIE%20Team--2021|The IMBIE Team, 2021]] ) leads to substantial observed surface lowering (e.g., [[#Schröder--2019|Schröder et al., 2019]] ; [[#Shepherd--2019|Shepherd et al., 2019]] ), particularly in coastal regions (Figure 9.18). Recent studies using satellite altimetry ( [[#Schröder--2019|Schröder et al., 2019]] ) and the input-output method ( [[#Rignot--2019|Rignot et al., 2019]] ) consistently show mass loss in these coastal regions since the late 1970s (Figure 9.16). Because of consistent multiple lines of evidence, there is ''high confidence'' in mass loss of the Totten Glacier in East Antarctica ( [[#Miles--2013|Miles et al., 2013]] ; X. [[#Li--2016|]] [[#Li--2016|Li et al., 2016]] ; [[#Mohajerani--2018|Mohajerani et al., 2018]] ; [[#Rignot--2019|Rignot et al., 2019]] ; [[#Schröder--2019|Schröder et al., 2019]] ; [[#Shepherd--2019|Shepherd et al., 2019]] ) since about 2000, dominated by changes in coastal ice dynamics (X. [[#Li--2016|]] [[#Li--2016|Li et al., 2016]] ). It is currently unclear whether mass loss of the EAIS over the last three decades has been significant ( [[#Rignot--2019|Rignot et al., 2019]] ) or, at 5 ± 46 Gt yr <sup>–1</sup> between 1992 and 2017, essentially zero within uncertainties ( [[#The%20IMBIE%20Team--2018|The IMBIE Team, 2018]] ). In summary, WAIS losses, through acceleration, retreat and thinning of the principal outlet glaciers, dominated the AIS mass losses over the last decades ( ''very high confidence'' ) and there is ''high confidence'' that this is the case since the late 1970s. Furthermore, parts of the EAIS have lost mass in the last two decades ( ''high confidence'' ).

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[[File:f9142cf45da1a832c41b4ffe2541a9e1 IPCC_AR6_WGI_Figure_9_18.png]]

'''Figure 9.18''' '''|''' '''Antarctic Ice Sheet cumulative mass change and equivalent sea level contribution. (a)''' A p-box ( [[#9.6.3.2|Section 9.6.3.2]] ) based estimate of the range of values of paleo Antarctic ice sheet mass and sea level equivalents relative to present day and the median over all central estimates ( [[#Bamber--2009|Bamber et al., 2009]] ; [[#Argus--2010|Argus and Peltier, 2010]] ; [[#Dolan--2011|Dolan et al., 2011]] ; [[#Mackintosh--2011|Mackintosh et al., 2011]] ; [[#Golledge--2012|Golledge et al., 2012]] , 2013, 2014, 2015, 2017b; K.G. [[#Miller--2012|]] [[#Miller--2012|Miller et al., 2012]] ; [[#Whitehouse--2012|Whitehouse et al., 2012]] ; [[#Ivins--2013|Ivins et al., 2013]] ; [[#Argus--2014|Argus et al., 2014]] ; [[#Briggs--2014|Briggs et al., 2014]] ; [[#Maris--2014|Maris et al., 2014]] ; [[#de%20Boer--2015|de Boer et al., 2015]] , 2017; [[#Dutton--2015|Dutton et al., 2015]] ; [[#Pollard--2015|Pollard et al., 2015]] ; [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#Gasson--2016|Gasson et al., 2016]] ; [[#Goelzer--2016|Goelzer et al., 2016]] ; [[#Yan--2016|Yan et al., 2016]] ; [[#Kopp--2017|Kopp et al., 2017]] ; [[#Simms--2019|Simms et al., 2019]] ); '''(b left)''' cumulative mass loss (and sea level equivalent) since 2015, with satellite observations shown from 1993 ( [[#Bamber--2018a|Bamber et al., 2018a]] ; [[#The%20IMBIE%20Team--2018|The IMBIE Team, 2018]] ; [[#WCRP%20Global%20Sea%20Level%20Budget%20Group--2018|WCRP Global Sea Level Budget Group, 2018]] ) and observations from 1979 ( [[#Rignot--2019|Rignot et al., 2019]] ), and projections from Ice Sheet Model Intercomparison Project for CMIP6 (ISMIP6) to 2100 under RCP8.5/SSP5-8.5 and RCP2.6/SSP1-2.6 scenarios (thin lines from [[#Seroussi--2020|Seroussi et al., 2020]] ; [[#Edwards--2021|Edwards et al., 2021]] ; [[#Payne--2021|Payne et al., 2021]] ) and ISMIP6 emulator under SSP5-8.5 and SSP1-2.6 to 2100 (shades and bold line; [[#Edwards--2021|Edwards et al., 2021]] ) ; (b, right) 17th–83rd, 5th–95th percentile ranges for ISMIP6, ISMIP6 emulator, and LARMIP-2 including surface mass balance (SMB) at 2100. (c – e) Schematic interpretations of individual reconstructions ( [[#Anderson--2002|Anderson et al., 2002]] ; [[#Bentley--2014|Bentley et al., 2014]] ; [[#de%20Boer--2015|de Boer et al., 2015]] ; [[#Goelzer--2016|Goelzer et al., 2016]] ) of the spatial extent of the Antarctic Ice Sheet are shown for the: '''(c)''' mid-Pliocene Warm Period, '''(d)''' Last Interglacial; and '''(e)''' Last Glacial Maximum ( [[#Fretwell--2013|Fretwell et al., 2013]] ): grey shading shows extent of grounded ice. (f – g) Maps of mean elevation changes '''(f)''' 1978–2017 derived from multi-mission satellite altimetry ( [[#Schröder--2019|Schröder et al., 2019]] ) and '''(g)''' ISMIP6: 2061–2100 projected changes for an ensemble using the Norwegian Climate Center’s Earth System Model (NorESM1-M) climate model under the RCP8.5 scenario ( [[#Seroussi--2020|Seroussi et al., 2020]] ). Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

As stated in SROCC, snowfall and glacier flow are the largest components determining AIS mass changes, with glacier flow acceleration (dynamic thinning) on the WAIS and the Antarctic Peninsula driving total loss trends in recent decades ( ''very high confidence'' ), and a partial offset of the dominating dynamic-thinning losses by increased snowfall ( ''high confidence'' ). The SROCC attributed ''medium confidence'' to estimates of 20th-century snowfall increases equivalent to a sea level change of –7.7 ± 4.0 mm on the EAIS, and –2.8 ± 1.7 mm on the WAIS, respectively ( [[#Medley--2019|Medley and Thomas, 2019]] ). Loss of buttressing, which can be caused by ice-shelf thinning, gradual ice-shelf front retreat or ice-shelf disintegration, has been linked to instantaneous ice velocity increases, and thus dynamic thinning, since the early 1990s. This link is clearly evident in the Amundsen and, to a lesser degree, Bellingshausen sectors ( [[#Gudmundsson--2019|Gudmundsson et al., 2019]] ), where passive shelf ice (ice that can be removed without major effects on the ice-shelf dynamics) is very limited or absent ( [[#Fürst--2016|Fürst et al., 2016]] ). Surface mass balance (SMB) changes, dominated by snowfall, exhibit strong regional and temporal variability, for example with multi-decadal increases in the Antarctic Peninsula inferred since the 1930s ( [[#Medley--2019|Medley and Thomas, 2019]] ), and dominate the interannual to decadal variability of the AIS mass balance ( [[#Rignot--2019|Rignot et al., 2019]] ). However, no significant continent-wide SMB trend is inferredsince 1979 ( [[#The%20IMBIE%20Team--2018|The IMBIE Team, 2018]] ; [[#Medley--2019|Medley and Thomas, 2019]] ; regional changes of Antarctic SMB are assessed further in [[IPCC:Wg1:Chapter:Atlas|Atlas]] [[IPCC:Wg1:Chapter:Chapter-11#11.1|Section 11.1]] ). In summary, there is ''very high confidence'' that the observed AIS mass loss since the early 1990s is primarily linked to ice-shelf changes.

The SROCC stated with ''high confidence'' that melting of ice shelves by warm ocean waters, leading to reduction of ice-shelf buttressing, has driven the observed ongoing thinning of major WAIS outlet glaciers. Since SROCC, digitized radar measurements have shown that the eastern ice shelf of Thwaites Glacier in the Amundsen Sea Embayment thinned between 10 and 33% during the three decades after 1978 ( [[#Schroeder--2019|Schroeder et al., 2019]] ), and the role of basal ice-shelf melting has been emphasized ( [[#Smith--2020|Smith et al., 2020]] ). Strong surface meltwater production has been noted as a precursor of ice-shelf disintegration in and since SROCC ( [[#Bell--2018|Bell et al., 2018]] ), and recent work placed strong meltwater production events ( [[#Lenaerts--2017|Lenaerts et al., 2017]] ; [[#Nicolas--2017|Nicolas et al., 2017]] ; [[#Wille--2019|Wille et al., 2019]] ) and seasons ( [[#Robel--2019|Robel and Banwell, 2019]] ) in this context. Antarctic ice-shelf basal meltwater flux varied between about 1100 ± 150 Gt yr <sup>–1</sup> in the mid-1990s and about 1570 ± 140 Gt yr <sup>–1</sup> in the late 2000s before decreasing to 1160 ± 150 Gt yr <sup>–1</sup> in 2018, and basal melt rates strongly vary with geographical position and depth, as a function of the surrounding water temperature ( [[#Adusumilli--2020|Adusumilli et al., 2020]] ). [[#9.2.2.3|Section 9.2.2.3]] assesses that the intrusion of warm Circumpolar Deep Water (CDW), which has warmed and shoaled since the 1980s, has been at least partially controlled by forcing with significant decadal variability ''. Limited evidence'' suggests that, beyond strong internal decadal wind variability, increased greenhouse gas forcing has slightly modified the mean local winds between 1920 and 2018, facilitating the intrusion of CDW heat on the Amundsen-Bellingshausen continental shelf, and increased ice shelf melt ( [[#9.2.2.3|Section 9.2.2.3]] ). However, theoretical understanding is still incomplete and in situ measurements within the ice–ocean boundary layer are sparse ( [[#Wåhlin--2020|Wåhlin et al., 2020]] ). Modelling, and therefore attribution of ice shelf basal melt, remains challenging because of insufficient process understanding, required spatial resolution, the paucity of in situ observations ( [[#Dinniman--2016|Dinniman et al., 2016]] ; [[#Asay-Davis--2017|Asay-Davis et al., 2017]] ; [[#Turner--2017|Turner et al., 2017]] ), and uncertainties of bathymetric datasets under ice-shelf cavities ( [[#Goldberg--2019|Goldberg et al., 2019]] , 2020; [[#Morlighem--2020|Morlighem et al., 2020]] ). In summary, ice-shelf thinning, mainly driven by basal melt, is widespread around the Antarctic coast and particularly strong around the WAIS ( ''high confidence'' ), although basal melt rates show substantial spatio-temporal variability.

Satellite observations suggest that changes in sea ice coverage and thickness can modulate iceberg calving, ice shelf flow and glacier terminus position around Antarctica ( [[#Miles--2013|Miles et al., 2013]] , 2016, 2017; [[#Massom--2015|Massom et al., 2015]] ; [[#Greene--2018|Greene et al., 2018]] ; [[#Bevan--2019|Bevan et al., 2019]] ), either through mechanical coupling or via changes to ocean stratification, influencing basal melting. A combined observational and modelling study ( [[#Massom--2018|Massom et al., 2018]] ) showed that regional loss of a protective sea ice buffer played a role in the rapid disintegration events of the Larsen A and B and Wilkins ice shelves in the Antarctic Peninsula between 1995 and 2009, by exposing damaged (rifted) outer ice shelf margins to enhanced flexure by storm-generated ocean swells. One observational study ( [[#Sun--2019|Sun et al., 2019]] ) suggests that the absence of sea ice in front of ice shelves, which leads to strengthened topographic waves, favours higher ice-shelf basal melt rates by increasing the baroclinic (depth varying) ocean heat flux which can enter the cavity ( [[#Wåhlin--2020|Wåhlin et al., 2020]] ). Paleo evidence for sea ice control on ice sheets is lacking, but geologic evidence shows a concordance between periods of ice-sheet growth and the expansion of sea ice ( [[#Patterson--2014|Patterson et al., 2014]] ; [[#Levy--2019|Levy et al., 2019]] ), both being favoured by reduced sea surface temperatures. Modelling confirms that sea ice controls the strength of ice mélange ( [[#Robel--2017|Robel, 2017]] ; [[#Schlemm--2021|Schlemm and Levermann, 2021]] ) and thus influences ice-shelf flexure and calving rates and stability of floating ice margins, but one model shows this had negligible effect on AIS retreat rates during past warm periods ( [[#Pollard--2018|Pollard et al., 2018]] ). Loss of ice-shelf-proximal sea ice is also associated with increased solar heating of surface waters and increased sub-shelf melting ( [[#Bendtsen--2017|Bendtsen et al., 2017]] ; [[#Stewart--2019|Stewart et al., 2019]] ). In summary, although in some cases sea ice decrease and glacier and ice-shelf flow and terminus position changes can have the same common cause, there is ''medium confidence'' that sea ice decrease ultimately favours the mass loss of nearby ice shelves through a variety of processes.

The SROCC stated with ''high confidence'' that ice-shelf disintegration has driven dynamic thinning in the northern Antarctic Peninsula over recent decades, and expressed ''high confidence'' in current ongoing mass loss from glaciers that fed now-disintegrated ice shelves. However, the mass loss rate has decreased in the 20 years since the immediate speed-up following ice-shelf disintegration in 1995 and 2002. Observed flow speed of these tributary glaciers is still 26% higher than before the ice shelf disintegration ( [[#Seehaus--2018|Seehaus et al., 2018]] ). Conversely, one study interpreted the increased flow speed of the Scar Inlet Ice Shelf’s tributary glaciers as a sign of evolving instability of the currently intact ice shelf ( [[#Qiao--2020|Qiao et al., 2020]] ).

Ongoing grounding line retreat, indicating dynamic thinning, is observed with ''high confidence'' in many areas of Antarctica, and particularly on the WAIS, with the highest rates being in the Amundsen and Bellingshausen Sea areas, and around Totten Glacier in East Antarctica, as stated in SROCC. Research published since SROCC has evidenced grounding line retreat of the West Antarctic Berry Glacier on the Getz Coast ( [[#Millan--2020|Millan et al., 2020]] ) and on the East Antarctic Denman Glacier ( [[#Brancato--2020|Brancato et al., 2020]] ), both since 1996. Furthermore observed grounding line retreat in excess of 1.5 km between 2003 and 2015 has been reported for parts of Marie Byrd Land ( [[#Christie--2018|Christie et al., 2018]] ). In summary, there is ''high confidence'' that grounding lines of marine-terminating glaciers are currently retreating in many areas around Antarctica, particularly around the WAIS, and additional areas of grounding line retreat have been evidenced since SROCC.

The SROCC stated with ''medium confidence'' that sustained mass losses of several major glaciers in the Amundsen Sea Embayment (ASE) are compatible with the onset of marine ice sheet instability (MISI). However, whether unstable WAIS retreat had begun, or was imminent, remained a critical uncertainty. New publications since SROCC have not substantially clarified this question. One study that combined satellite measurements with a numerical model and prescribed ice-shelf thinning ( [[#Gudmundsson--2019|Gudmundsson et al., 2019]] ) suggests that MISI is not required to explain the observed current mass loss rates of the WAIS, because they are consistent with external climate drivers. Furthermore, the fast grounding line retreat of the Pine Island Glacier in the ASE, which was triggered in the 1940s ( [[#Smith--2017|Smith et al., 2017]] ), observed after 1992 ( [[#Rignot--2014|Rignot et al., 2014]] ) and previously interpreted as a sign of MISI ( [[#Favier--2014|Favier et al., 2014]] ), seems to have stabilized recently ( [[#Milillo--2017|Milillo et al., 2017]] ; [[#Konrad--2018|Konrad et al., 2018]] ), and its current flow patterns do not suggest ongoing or imminent MISI ( [[#Bamber--2020|Bamber and Dawson, 2020]] ). However, sustained fast grounding line retreat has been observed for the Smith Glacier in the ASE ( [[#Scheuchl--2016|Scheuchl et al., 2016]] ), and an analysis of flow patterns and grounding line retreat of the ASE Thwaites Glacier between 1992 and 2017 ( [[#Milillo--2019|Milillo et al., 2019]] ) showed sustained, albeit spatially heterogeneous, grounding line retreat, highlighting ice–ocean interactions that lead to increased basal melt. In addition, Denman Glacier in East Antarctica was shown to hold potential for unstable retreat ( [[#Brancato--2020|Brancato et al., 2020]] ). In summary, the observed evolution of the ASE glaciers is compatible with, but not unequivocally indicating an ongoing MISI ( ''medium confidence'' ).

The SROCC reported ''limited evidence'' and ''medium agreement'' for anthropogenic forcing of the observed AIS mass balance changes. As stated in [[IPCC:Wg1:Chapter:Chapter-3#3.4.3.2|Section 3.4.3.2]] , there remains ''low confidence'' in attributing the causes of the observed mass of loss from the AIS since 1993, in spite of some additional process-based evidence to support attribution to anthropogenic forcing.

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==== 9.4.2.2 Model Evaluation ====

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The AR5 ( [[#Church--2013b|Church et al., 2013b]] ; [[#Flato--2013|Flato et al., 2013]] ) stated that regional climate models and global models with bias-corrected SST and sea ice concentration tended to produce more accurate simulations of Antarctic SMB than coupled climate models. It also noted strong climate model temperature biases over the Antarctic, though the latter may reflect known biases in the reanalysis used ( [[#Fréville--2014|Fréville et al., 2014]] ). Section Atlas.11.1 assesses that there is ''medium confidence'' in the capacity of climate models to simulate Antarctic climatology and SMB changes.

( [[#9.2.3.2|Section 9.2.3.2]] assesses that there is ''low confidence'' in simulations of Southern Ocean temperature. Few ocean models resolve ice-shelf cavities, and biases in present-day melt rates can be substantial in some sectors, including the key region of the Amundsen Sea (e.g., an exception is the FESOM simulation in Figure 9.19 includes ice-shelf cavities and simulates ice-shelf basal melting and refreezing) ( [[#Naughten--2018|Naughten et al., 2018]] ). An increasing number of observational studies from which basal melt rates are calculated ( [[#Huhn--2018|Huhn et al., 2018]] ; [[#Adusumilli--2020|Adusumilli et al., 2020]] ; [[#Das--2020|Das et al., 2020]] ; [[#Hirano--2020|Hirano et al., 2020]] ; [[#Stevens--2020|Stevens et al., 2020]] ), combined with improved understanding of influences specific to water-masses and modes of melting or dissolving ( [[#Silvano--2018|Silvano et al., 2018]] ; [[#Adusumilli--2020|Adusumilli et al., 2020]] ; [[#Malyarenko--2020|Malyarenko et al., 2020]] ; [[#Wåhlin--2020|Wåhlin et al., 2020]] ), may help to refine these models in the future. However, given the limited number of available models and their biases, there is currently ''low confidence'' in the sub-shelf melt rates simulated by ocean models.

Improvements in the representation of grounding line evolution in ice-sheet models since AR5 (such as sub-grid schemes for basal friction and ice-shelf melt, and local grid refinement) means that most of the model simulations presented in SROCC were dominated by physical processes. Since then, these advances have been applied in several model intercomparison projects – such as ISMIP6 and LARMIP-2 (see Box 9.3); MISMIP+ (Cornford et al. 2020); and ABUMIP (Sun et al. 2020). All models participating in ISMIP6 and LARMIP-2 simulate ice-shelf and grounding-line evolution, and include sub-shelf melt parametrization, which was not the case in the Sea-level Response to Ice Sheet Evolution (SeaRISE) project intercomparison ( [[#Bindschadler--2013|Bindschadler et al., 2013]] ; [[#Nowicki--2013|Nowicki et al., 2013]] ). Simulations of grounding line evolution ( [[#Seroussi--2017|Seroussi et al., 2017]] , 2020) have benefitted from improved bedrock topography ( [[#Morlighem--2020|Morlighem et al., 2020]] ). Treatment of sub-shelf melting, however, remains one of the causes of large differences in AIS models, particularly for partially floating grid cells in models with coarse resolution ( [[#Levermann--2020|Levermann et al., 2020]] ; [[#Edwards--2021|Edwards et al., 2021]] ). Due to the limitations in resolving cavities in ocean models, as described above, basal melt rates are generally parameterized at the ice shelf base, based on ocean model simulations of temperatures and salinity instead ( [[#Nowicki--2020b|Nowicki et al., 2020b]] ; [[#Seroussi--2020|Seroussi et al., 2020]] ). While this has the advantage of connecting melt rates to emissions scenarios, a large variety of melt parametrizations exist ( [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#Lazeroms--2018|Lazeroms et al., 2018]] ; [[#Reese--2018|Reese et al., 2018]] ; [[#Hoffman--2019|Hoffman et al., 2019]] ; [[#Pelle--2019|Pelle et al., 2019]] ; [[#Jourdain--2020|Jourdain et al., 2020]] ), and there is ''low agreement'' due to limited observational constraints (ocean temperature, salinity, velocity, and ice shelf draft)( [[#Jourdain--2020|Jourdain et al., 2020]] ), uncertainty in the physics of parametrized processes, missing processes (e.g., tides), and uncertainty in the treatment of ice-sheet–climate feedbacks ( [[#Donat-Magnin--2017|Donat-Magnin et al., 2017]] ; [[#Bronselaer--2018|Bronselaer et al., 2018]] ; [[#Golledge--2019|Golledge et al., 2019]] ). Parametrizations are usually calibrated to present-day melt rates, but can respond differently to projected ocean warming ( [[#Favier--2019|Favier et al., 2019]] ; [[#Jourdain--2020|Jourdain et al., 2020]] ). Two different calibrations were used in ISMIP6 (Box 9.3; [[#Jourdain--2020|Jourdain et al., 2020]] ; [[#Nowicki--2020b|Nowicki et al., 2020b]] ): one reproducing melt rates averaged around the whole continent (MeanAnt: Figure 9.19), and the other reproducing melt rates near the grounding line of Pine Island Glacier (PIGL; see Figure 9.19), leading to large differences in melt rates. Evaluation with observations and two cavity-resolving models suggests that the MeanAnt parametrization better reproduces observed melt rates and projected increases in both the warm Amundsen Sea Embayment and cold Ronne-Filchner shelf cavity, as well as total Antarctic melting ( [[#Jourdain--2020|Jourdain et al., 2020]] ). The PIGL calibration represents the upper end for increased basal melt sensitivity that would be caused by continent-wide changes to ocean water properties and circulation under strong future forcing ( [[#Jourdain--2020|Jourdain et al., 2020]] ). The basal sliding law also has a strong influence on grounding line retreat and glacier acceleration in response to perturbations, and varies spatially ( [[#Sun--2020|Sun et al., 2020]] ). Sliding laws ( [[#Joughin--2019|Joughin et al., 2019]] ) can only be constrained with observations in regions experiencing significant change, and with sufficiently long observational records.

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[[File:d2fb69c7e007aee6fe14976159bbf4c8 IPCC_AR6_WGI_Figure_9_19.png]]

'''Figure 9.19''' '''|''' '''Ice-shelf basal melt rates for present-day (upper panels) and changes from present-day to the end of the 21st century under the RCP8.5 scenario (lower panels).''' Present-day melt rates were estimated through: the input-output method constrained by satellite observations and atmosphere/snow simulations ( [[#Rignot--2013|Rignot et al., 2013]] ) and representative of 2003–2008 (upper left); the Ice Sheet Model Intercomparison Project for CMIP6 (ISMIP6) non-local-PIGL parametrization constrained by observation-based ocean properties ( [[#Jourdain--2020|Jourdain et al., 2020]] ) and representative of 1995–2014 (upper centre); the Finite Element Sea ice/Ice Shelf Ocean Model (FESOM) simulation over 2006–2015, forced by atmospheric conditions from a Coupled Model Intercomparison Project Phase 5 (CMIP5) multi-model mean (MMM) under the RCP8.5 scenario ( [[#Naughten--2018|Naughten et al., 2018]] ) (upper right). Future anomalies are calculated as 2081–2100 minus present-day using the ISMIP6 non-local-MeanAnt and non-local-PIGL parametrizations ( [[#Jourdain--2020|Jourdain et al., 2020]] ) (lower left and centre, respectively) based on projections from the Norwegian Climate Center’s Earth System Model (NorESM1-M) CMIP5 model, and the FESOM-MMM projection (lower right). Note the symmetric-log colour bar (linear around zero, logarithmic for stronger negative and positive values). Inset highlights the Amundsen Sea Region. Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

The SROCC noted that AIS simulations are increasingly evaluated or formally calibrated with modern observations and/or paleodata – to obtain more realistic initial conditions (ice-sheet geometry, velocity and forcing) and to constrain uncertainty in probabilistic projections. This trend continues ( [[#Nias--2019|Nias et al., 2019]] ; [[#Gilford--2020|Gilford et al., 2020]] ; [[#Hamlington--2020b|Hamlington et al., 2020b]] ; [[#Wernecke--2020|Wernecke et al., 2020]] ). However, while the large-scale characteristics of the initial ice-sheet state have improved significantly (Box 9.3), capturing the smaller-scale rates of change, including mass trends, remains challenging for many models ( [[#Goldberg--2015|Goldberg et al., 2015]] ; [[#Reese--2020|Reese et al., 2020]] ; [[#Seroussi--2020|Seroussi et al., 2020]] ; [[#Siegert--2020|Siegert et al., 2020]] ). This increases uncertainty in projections, especially for the 21st century ( [[#9.4.2.5|Section 9.4.2.5]] ). However, uncertainties in ice-sheet model simulations have been much better quantified since AR5, through model intercomparison projects (in particular, ISMIP6 and LARMIP-2; see Box 9.3), perturbed parameter ensembles, and increasing use of statistical emulation ( [[#Gilford--2020|Gilford et al., 2020]] ; [[#Levermann--2020|Levermann et al., 2020]] ; [[#Wernecke--2020|Wernecke et al., 2020]] ; [[#DeConto--2021|DeConto et al., 2021]] ; [[#Edwards--2021|Edwards et al., 2021]] ) to better sample the parameter space. By exploring uncertainties more fully, these methods have the potential to identify better simulations of the historical period.

An important difficulty is how to evaluate simulations of processes that are: not currently observed; or rare; or indirectly deduced – in particular, the ice-shelf disintegrations and cliff failures that would drive the proposed marine ice cliff instability (MICI; [[#9.4.2.4|Section 9.4.2.4]] and Box 9.4; [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#DeConto--2021|DeConto et al., 2021]] ). Models of ice-cliff failure can only be indirectly and partially evaluated, using existing (i.e., static) cliffs and laboratory experiments ( [[#Clerc--2019|Clerc et al., 2019]] ). The SROCC stated that there was ''low agreement'' on the exact MICI mechanism and ''limited evidence'' of its occurrence in the present or the past, and that the validity of MICI remains unproven. Only one ice-sheet model represents MICI ( [[#Pollard--2015|Pollard et al., 2015]] ; [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#DeConto--2021|DeConto et al., 2021]] ). The mechanism has not been found to be essential for reproducing Mid Pliocene Warm Period and Last Interglacial reconstructions or satellite observations, though Last Interglacial data slightly favours it in this model ( [[#Edwards--2019|Edwards et al., 2019]] ; [[#Gilford--2020|Gilford et al., 2020]] ; [[#DeConto--2021|DeConto et al., 2021]] ).

In summary, there is now ''medium confidence'' in many ice-sheet processes in ice-sheet models, including grounding line evolution. However, there remains ''low confidence'' in the ocean forcing affecting the basal melt rates, and ''low confidence'' in simulating mechanisms that have the potential to cause widespread, sustained and very rapid ice loss from Antarctica through MICI.

<div id="9.4.2.3" class="h3-container"></div>

<span id="drivers-of-future-antarctic-ice-sheet-change"></span>
==== 9.4.2.3 Drivers of Future Antarctic Ice Sheet Change ====

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<div id="9.4.2.3.1" class="h4-container"></div>

<span id="surface-mass-balance"></span>
===== 9.4.2.3.1 Surface mass balance =====

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The AR5 projected a negative contribution from Antarctic surface mass balance (SMB) changes to sea level over the 21st century (i.e., mitigating sea level rise), due to increased snowfall associated with warmer air temperatures. Sensitivity of SMB to Antarctic surface air temperature change varied from 3.7 to 7% °C <sup>–1</sup> , and the sea level projections assumed a sensitivity of 5.1 ± 1.5% °C <sup>–1</sup> from CMIP3 era models ( [[#Gregory--2006|Gregory and Huybrechts, 2006]] ) to estimate SMB changes from Antarctic temperatures in the CMIP5 ensemble. Since the AR5, analyses of CMIP5 and CMIP6 models have found Antarctic temperature sensitivity for accumulation (precipitation minus sublimation) of 3.5 to 8.7% °C <sup>–1</sup> ( [[#Frieler--2015|Frieler et al., 2015]] ), for SMB of 6.0 to 9.9% °C <sup>–1</sup> ( [[#Previdi--2016|Previdi and Polvani, 2016]] ) and for precipitation of around 4 to 9% °C <sup>–1</sup> (±1 standard deviation ranges; [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ). An accumulation sensitivity estimate derived from ice core data lies in the middle of the range, around 6% °C <sup>–1</sup> ( [[#Frieler--2015|Frieler et al., 2015]] ). These are consistent, within uncertainties, with each other and AR5, under the approximation that SMB is dominated by snowfall.

The AR5 found that the median and ''likely'' sea level contributions due to SMB from 1986–2005 to 2100 were –0.05 (–0.09 to –0.02) m under RCP8.5 and –0.02 (–0.05 to 0.00) m under RCP2.6. The SROCC did not present a separate SMB contribution, instead showing total Antarctic projections derived from ice-sheet models ( [[#9.4.2.5|Section 9.4.2.5]] ). Projections of the SMB contribution to sea level tend to be slightly more negative since AR5, due at least in part to the higher range in equilibrium climate sensitivity values in CMIP6 ( [[#Payne--2021|Payne et al., 2021]] ). Mean and ±1 standard deviation ranges for grounded Antarctic Ice Sheet SMB changes from 2000 to 2100 computed from CMIP5 models are –0.08 (–0.13 to –0.04) m sea level equivalent (SLE) for RCP8.5 and, similarly for CMIP6 models, are –0.07 (–0.11 to –0.03) m for SSP5-8.5 ( [[#Gorte--2020|Gorte et al., 2020]] ). The general circulation models (GCMs) used to drive ice-sheet models in ISMIP6 (Box 9.3) project mean grounded AIS SMB changes from 2005 to 2100 of –0.06 (range –0.08 to –0.03) m SLE under RCP8.5 for the six CMIP5 models ( [[#Seroussi--2020|Seroussi et al., 2020]] ) and –0.09 (range –0.10 to –0.07) m SLE under SSP5-8.5 for the four CMIP6 models, which have climate sensitivity values of 4.8°C –5.3°C ( [[#Payne--2021|Payne et al., 2021]] ). We apply the AR5 parametric AIS SMB model ( [[#9.6.3.2|Section 9.6.3.2]] ) to updated projections of global mean temperature from a two-layer energy budget emulator (Supplementary Material 7.SM.2), which gives a median –0.05 (5–95% range –0.07 to –0.02) m SLE for SSP5-8.5 ( [[#9.4.2.5|Section 9.4.2.5]] , Table 9.3), that is, similar to the AR5 assessment and slightly smaller than the CMIP6 estimate. This estimate is used to augment the LARMIP-2 dynamic projections (Box 9.3) in Sections 9.4.2.5 and 9.4.2.6. Overall, CMIP5 and CMIP6 GCM simulations of sea level fall by 2100 due to Antarctic SMB increases are around 2–4 cm greater than estimates derived with the statistical method used in AR5. Further details about projections of Antarctic temperature, precipitation and SMB are provided in Section Atlas.11.1.4, which assesses that, due to the challenges of model evaluation ( [[#9.4.2.2|Section 9.4.2.2]] ) and the possibility of increased meltwater runoff ( [[#Kittel--2021|Kittel et al., 2021]] ), there is only ''medium confidence'' that the future contribution of Antarctic SMB to sea level this century will be negative under all greenhouse gas emissions scenarios. Longer time scales are discussed in 9.4.2.6.

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<span id="sub-shelf-melting"></span>
===== 9.4.2.3.2 Sub-shelf melting =====

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The SROCC highlighted that an important ongoing deficiency in projections of Antarctic sub-shelf melting is the lack of ice–ocean coupling in most continental-scale studies. Increased basal melting is mainly caused by warmer CDW ( [[#9.2.2.3|Section 9.2.2.3]] ) on the continental shelves, and warming surface waters intruding under ice shelves ( [[#Naughten--2018|Naughten et al., 2018]] ). Predicting whether or not open ocean water masses will freely penetrate ice shelf cavities, or will be partially blocked by ocean density gradients, is complex ( [[#Wåhlin--2020|Wåhlin et al., 2020]] ); while melting related to CDW inflow is currently dominant in the Amundsen Sea Embayment, melt in other embayments is limited by deep inflows of high-salinity shelf water or seasonally warmed shallow incursions of Antarctic Surface Water ( [[#Stewart--2019|Stewart et al., 2019]] ; [[#Adusumilli--2020|Adusumilli et al., 2020]] ). There is little consensus regarding future change in CDW ( [[#9.2.2.3|Section 9.2.2.3]] ), and more generally ''low confidence'' in future change in the temperature of Antarctic ice-shelf cavities ( [[#9.2.3.2|Section 9.2.3.2]] ).

The response of sub-shelf melting to ocean warming is also poorly constrained. A key unknown is whether, and when, cold ice-shelf cavities might become more similar to the Amundsen Sea Embayment, not only in ocean temperature but also ice–ocean heat exchange, which depends on the cavity geometry and ocean circulation ( [[#Little--2009|Little et al., 2009]] ). Only two ocean models with ice-shelf cavities have been used to make sub-shelf basal melting projections for Special Report on Emissions Scenarios and Representative Concentration Pathway (RCP) scenarios ( [[#Hellmer--2012|Hellmer et al., 2012]] ; [[#Timmermann--2013|Timmermann and Hellmer, 2013]] ; [[#Timmermann--2017|Timmermann and Goeller, 2017]] ; [[#Naughten--2018|Naughten et al., 2018]] ). The FESOM simulation, forced by a CMIP5 multi-model mean under RCP8.5, projects a 90% increase in melting (Figure 9.19), although this could be overestimated due to an underestimation of present-day melt rates ( [[#9.4.2.2|Section 9.4.2.2]] ; [[#Naughten--2018|Naughten et al., 2018]] ). The temperature–melt relationship was parameterized by ISMIP6 in terms of heat exchange velocity in m a <sup>–1</sup> , and by LARMIP-2 as basal melt sensitivity in m a <sup>–1</sup> °C <sup>–1</sup> (Box 9.3; [[#Jourdain--2020|Jourdain et al., 2020]] ; [[#Levermann--2020|Levermann et al., 2020]] ; [[#Reese--2020|Reese et al., 2020]] ), and both vary widely around the continent, depending on cavity type. Median values of ISMIP6 heat exchange velocity vary by a factor of 5–10 when calibrating to either mean Antarctic or high Pine Island Glacier observed melt rates ( [[#9.4.2.2|Section 9.4.2.2]] ; Box 9.3; [[#Jourdain--2020|Jourdain et al., 2020]] ). Basal melt sensitivities near the grounding line estimated by [[#Reese--2020|Reese et al. (2020)]] with a box model of ocean overturning range from 3.9 m a <sup>–1</sup> °C <sup>–1</sup> for the Weddell Sea to 10.5 m a <sup>–1</sup> °C <sup>–1</sup> for the Amundsen Sea region, with a continental mean of 5.3 m a <sup>–1</sup> °C <sup>–1</sup> . Similarly high Amundsen Sea sensitivities are estimated in coupled ice–ocean simulations of Thwaites Glacier (mean 9.4 m a <sup>–1</sup> °C <sup>–1</sup> ; range 6–16 m a <sup>–1</sup> °C <sup>–1</sup> ) ( [[#Seroussi--2017|Seroussi et al., 2017]] ). These large variations lead to large differences in basal melt rates and projected sea level contributions when applied to the whole ice sheet in ISMIP6 and LARMIP-2 (Box 9.3). Projections of melt rates from the two ISMIP6 calibrations are higher than those from FESOM, driven by a CMIP5 multi-model mean (Figure 9.19; [[#Jourdain--2020|Jourdain et al., 2020]] ). The ISMIP6 ensemble mostly uses the mean Antarctic calibration, but includes some simulations with the Pine Island Glacier calibration, and the ISMIP6 emulator samples more of these higher values; LARMIP-2 uses basal melt sensitivities (7–16 m a <sup>–1</sup> °C <sup>–1</sup> ) consistent with estimates for the Amundsen Sea Embayment. Due to the limited availability of cavity-resolving ocean models, and the wide regional variation in estimates of basal melt sensitivity to ocean temperature, there is only ''low confidence'' in projected future sub-ice-shelf melt rates. The impact of this uncertainty on AIS model projections to 2100 is discussed in [[#9.4.2.5|Section 9.4.2.5]] .

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<span id="ice-shelf-disintegration"></span>
===== 9.4.2.3.3 Ice-shelf disintegration =====

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Antarctic ice shelves modulate grounded ice flow through buttressing, so their weakening or disintegration is crucial for the timing and magnitude of ice loss and onset of instabilities ( [[#9.4.2.4|Section 9.4.2.4]] ; Box 9.4). Projections of ice-shelf disintegration are uncertain in terms of atmospheric warming and the response of the shelf surface – that is, surface melting, and whether shelves then disintegrate due to hydrofracturing and flexing, or are resilient through refreezing or drainage ( [[#Bell--2018|Bell et al., 2018]] ). The SROCC stated it is not expected that widespread ice-shelf loss will occur before the end of the 21st century, but this was based on only one study, using a regional climate model forced by five GCMs ( [[#Trusel--2015|Trusel et al., 2015]] ), so there was ''low confidence'' in this assessment. The study of [[#DeConto--2016|DeConto and Pollard (2016)]] projected the appearance of extensive surface meltwater several decades earlier than [[#Trusel--2015|Trusel et al. (2015)]] and was therefore assessed to be too uncertain to include in SROCC projections of the AIS.

Since SROCC, further studies have highlighted the modelling uncertainties in this area. Coastal surface air temperature projections in CMIP6 models show large inter-model differences driven by sea ice retreat and exhibit more warming relative to global mean temperature under low emissions than high, due to delayed response of the Southern Ocean to stabilized emissions and stratospheric ozone recovery ( [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ). The updated study of [[#DeConto--2021|DeConto et al. (2021)]] includes improvements to the climate simulations relative to those in [[#DeConto--2016|DeConto and Pollard (2016)]] , and the resulting surface meltwater projections are now consistent with [[#Trusel--2015|Trusel et al. (2015)]] . However, the net effect of meltwater feedbacks on ice shelves is uncertain. Ice discharge is expected to lead to surface ocean and atmosphere cooling: this increases ocean stratification and sub-shelf melting, but also reduces ice-shelf surface melting and delays hydrofracturing ( [[#Golledge--2019|Golledge et al., 2019]] ; [[#Sadai--2020|Sadai et al., 2020]] ; [[#DeConto--2021|DeConto et al., 2021]] ). The new studies are insufficient to change SROCC’s ''low confidence'' assessment on ice-shelf loss. The consequence of this uncertainty on projections is discussed in [[#9.4.2.5|Section 9.4.2.5]] and Box 9.4.

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<span id="ice-sheet-instabilities"></span>
==== 9.4.2.4 Ice-sheet Instabilities ====

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A major uncertainty in future Antarctic mass losses is the possibility of rapid and/or irreversible ice losses through instability of marine parts of the ice sheet, via the proposed mechanisms of marine ice sheet instability (MISI) and marine ice cliff instability (MICI), and whether these processes will lead to a collapse of the West Antarctic Ice Sheet (WAIS).

MISI is a proposed self-reinforcing mechanism within marine ice sheets that lie on a bed that slopes down towards the interior of the ice sheet, whereby, in the absence of ice-shelf buttressing, the position of the grounding line is inherently unstable until reaching an upward sloping bed. The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) noted advances in modelling MISI since AR5, but that ‘significant discrepancies’ remained in projections due to poor understanding of mechanisms, and lack of observational data to constrain the models. Since SROCC, modelling uncertainties have been more thoroughly explored, rather than constrained (compatibility of current observations in the Amundsen Sea Embayment with MISI is assessed in [[#9.4.2.1|Section 9.4.2.1]] ). Internal climate variability might either slow ( [[#Hoffman--2019|Hoffman et al., 2019]] ) or amplify ( [[#Robel--2019|Robel et al., 2019]] ) MISI, and stable grounding line positions can be reached on downward sloping beds if ice shelves provide buttressing ( [[#Sergienko--2019|Sergienko and Wingham, 2019]] ; [[#Cornford--2020|Cornford et al., 2020]] ). Ice-sheet model simulations that remove all Antarctic ice shelves (and prevent them from reforming) show 2–10 m SLE Antarctic mass loss after 500 years due to MISI, of which WAIS collapse contributes 2–5 m ( [[#Sun--2020|Sun et al., 2020]] ), with the majority of the mass loss in the first one to two centuries. Much of the multi-model variation is due to the sliding law ( [[#9.4.2.2|Section 9.4.2.2]] ). However, it is not expected that widespread ice-shelf loss will occur before the end of the 21st century ( [[#9.4.2.3|Section 9.4.2.3]] ; Box 9.4). A recent update of bed topography that unveiled large and overdeepened subglacial troughs in East Antarctica potentially vulnerable to MISI ( [[#Morlighem--2020|Morlighem et al., 2020]] ) has only been used by a few models ( [[#Seroussi--2020|Seroussi et al., 2020]] ; [[#Sun--2020|Sun et al., 2020]] ), so current projections could underestimate vulnerability in these regions. The sea level rise contribution of the AIS therefore crucially depends on the behaviour of individual ice shelves and outlet glacier systems and whether they enter MISI for a given level of warming (Box 9.4; [[#Pattyn--2020|Pattyn and Morlighem, 2020]] ). As for Antarctic simulations generally (Sections 9.4.2.2 and 9.4.2.3), there is ''medium confidence'' in simulating MISI but ''low confidence'' in projecting the sub-shelf melting and ice-shelf disintegration that drive it.

The SROCC noted ''limited evidence'' from geological records and ice-sheet modelling, suggesting that parts of the AIS experienced rapid (centennial) retreat ''likely'' due to MISI between 20,000 and 9,000 years ago, and also described more uncertain evidence for the Last Interglacial (LIG) and mid-Pliocene Warm Period (MPWP). Recent support for past MISI is provided by model simulations of the WAIS during the LIG ( [[#Clark--2020|Clark et al., 2020]] ), the British Ice Sheet during the last termination ( [[#Gandy--2018|Gandy et al., 2018]] ) and the Laurentide Ice Sheet during the Younger Dryas ( [[#Pico--2019|Pico et al., 2019]] ), which show progressive retreat despite declining temperatures, indicative of a true (ice dynamic) instability. Direct observational evidence of rapid paleo ice-sheet grounding line retreat is rare but, on the Larsen continental shelf, retreat rates of &gt;10 km yr <sup>–1</sup> during the deglaciation have been estimated ( [[#Dowdeswell--2020|Dowdeswell et al., 2020]] ). MISI has also been inferred from sedimentological evidence of ice loss from Wilkes Subglacial Basin, East Antarctica ( [[#Bertram--2018|Bertram et al., 2018]] ; [[#Wilson--2018|Wilson et al., 2018]] ; [[#Blackburn--2020|Blackburn et al., 2020]] ) but these reconstructions cannot unambiguously identify unstable from progressive retreat. Therefore, there is ''limited evidence'' to identify the operation of instability mechanisms such as MISI in paleo ice-sheet retreat.

The SROCC assessed that ice-sheet interactions with the solid Earth are not expected to substantially slow sea level rise from marine-based ice in Antarctica over the 21st century ( ''medium confidence'' ), but that these processes could become important on multi-century and longer time scales. More recent modelling of deglaciation of the Ross Embayment by [[#Lowry--2020|Lowry et al. (2020)]] is consistent with this assessment. However, new projections for Pine Island Glacier ( [[#Kachuck--2020|Kachuck et al., 2020]] ) support previous work ( [[#Barletta--2018|Barletta et al., 2018]] ) suggesting that lower mantle viscosity in this region leads to a negative feedback on decadal time scales. Grounding line stabilization by the solid Earth response may therefore occur over the 21st century in the Amundsen Sea Embayment, where most mass loss is occurring ( [[#9.4.2.1|Section 9.4.2.1]] ), but more generally occurs over multi-centennial to millennial time scales ( ''medium confidence'' ).

The MICI hypothesis describes rapid, unmitigated calving triggered by ice-shelf collapse ( [[#Pollard--2015|Pollard et al., 2015]] ). The SROCC noted that the MICI mechanism led one model ( [[#DeConto--2016|DeConto and Pollard, 2016]] ) to lose mass far more rapidly, but excluded the mechanism from its projections due to uncertainty in the timing of the ice-shelf disintegration ( [[#9.4.2.3|Section 9.4.2.3]] ). They stated that MICI could lead to sea level contributions beyond 2100 considerably higher than the ''likely'' range projected by other models. However, given the ''low agreement'' on the exact MICI mechanism and ''limited evidence'' of its occurrence in the present or the past ( [[#9.4.2.2|Section 9.4.2.2]] ), its potential to affect future sea level rise was very uncertain. Since SROCC, new simulations show later ice-shelf disintegration, in agreement with other models ( [[#9.4.2.3|Section 9.4.2.3]] ; [[#DeConto--2021|DeConto et al., 2021]] ), and therefore lower projections at 2100 ( [[#9.4.2.5|Section 9.4.2.5]] ). New theoretical evidence suggests that ice-cliff collapse may only occur after very rapid ice shelf disintegration caused by unusually high meltwater production ( [[#Clerc--2019|Clerc et al., 2019]] ; [[#Robel--2019|Robel and Banwell, 2019]] ), and that the subsequent rate of retreat depends on the terminus geometry ( [[#Bassis--2019|Bassis and Ultee, 2019]] ). As SROCC noted, only Crane Glacier on the Peninsula has shown retreat consistent with MICI, after the Larsen B ice shelf collapsed, and MICI-style behaviour at Jakobshavn and Helheim Glaciers in Greenland might not be representative of wider Antarctic glaciers. Observations from Greenland show that steep cliffs commonly evolve into short floating extensions, rather than collapsing catastrophically ( [[#Joughin--2020|Joughin et al., 2020]] ). As assessed in [[#9.4.2.2|Section 9.4.2.2]] and 9.4.2.3, there is therefore ''low confidence'' in simulating mechanisms that have the potential to cause widespread, sustained and very rapid ice loss from Antarctica this century through MICI, and ''low confidence'' in projecting the driver of ice-shelf disintegration.

In summary, poorly understood processes of instabilities, characterized by ''deep uncertainty'' , have the potential to strongly increase Antarctic mass loss under high greenhouse gas emissions on century-to-multicentury time scales (Box 9.4). These instabilities are therefore considered separately in assessments of the future contribution to global mean sea level (GMSL; Sections 9.4.2.5, 9.4.2.6, 9.6.3.2 and 9.6.3.5).

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<span id="projections-to-2100-1"></span>
==== 9.4.2.5 Projections to 2100 ====

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The AR5 assessed the median and ''likely'' (66–100% probability) sea level contributions of the AIS in 2100 relative to 1986–2005 to be 0.06 (–0.04 to +0.16) m SLE under RCP2.6 and 0.04 (–0.08 to +0.14) m SLE under RCP8.5 (Table 9.3; no change when using the AR6 baseline). The AR5 stated that only the collapse of the marine-based sectors of the AIS, if initiated, could cause GMSL to rise substantially above the ''likely'' range during the 21st century, with ''medium confidence'' that this would not exceed several tenths of a metre during this period. The assessment of the dynamical contribution had no dependence on emissions scenarios, due to the lack of literature, so the decrease in sea level contribution in the higher-emissions scenario was solely due to increased SMB ( [[#9.4.2.3|Section 9.4.2.3]] ). The SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ) assessed the total contribution based on five new ice-sheet modelling studies that incorporated marine ice-sheet dynamics, combining their estimates and interpreting the 5–95th percentile range of the resulting distribution as the ''likely'' range (17–83% probability interval, i.e., not open-ended as in the AR5). The median and ''likely'' range contributions by 2100 were 0.04 (0.01–0.11) m under RCP2.6 and 0.12 (0.03–0.28) m under RCP8.5 (Table 9.3). The positive scenario-dependence in SROCC – where increases in dynamic losses driven by ocean warming and ice-shelf disintegration under higher emissions ( [[#9.4.2.3|Section 9.4.2.3]] ) dominate over increases in SMB – arose from a combination of physical processes and model limitations. Modelling improvements in these studies included improved representations of grounding line response to drivers, more extensive exploration of uncertainties, and inclusion of a positive feedback of meltwater on climate ( [[#Golledge--2019|Golledge et al., 2019]] ). However, two of the projections did not include SMB changes that would offset dynamic losses ( [[#Levermann--2014|Levermann et al., 2014]] ; [[#Ritz--2015|Ritz et al., 2015]] ), and the scenario dependence may have been further amplified by highly sensitive sub-shelf melt parametrizations and use of simplified SMB schemes ( [[#Golledge--2015|Golledge et al., 2015]] , 2019; [[#Bulthuis--2019|Bulthuis et al., 2019]] ; [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ).

Since SROCC, new projections have arisen from multi-model intercomparison projects ISMIP6 and LARMIP-2 (Box 9.3) and one model that includes MICI ( [[#9.4.2.4|Section 9.4.2.4]] ; Table 9.3; [[#DeConto--2021|DeConto et al., 2021]] ). Corrections are added to allow comparison: all ISMIP6-derived projections have an estimate of the historical dynamical response to pre-2015 climate forcing added, which increases contributions (Box 9.3; Figure 9.18); the LARMIP-2 dynamic projections are combined with an estimate of SMB, which decreases contributions (Sections 9.4.2.3 and 9.6.3.2); and the ISMIP6 emulated and LARMIP-2 projections were re-estimate using the global surface air temperature distributions from the two-layer energy budget emulator described in Supplementary Material 7.SM.2. The majority of the new projections indicate that, under all emissions scenarios, the AIS will lose mass overall and contribute to sea level rise. Most thinning occurs in the Amundsen Sea sector in WAIS and Totten Glacier in EAIS (Figure 9.18). The most negative contribution is –0.02 m (5th percentile of ISMIP6 combined RCP8.5 and SSP5-8.5 projections after correction) and the largest contribution is 0.57 m SLE (95th percentile; [[#Levermann--2020|Levermann et al., 2020]] ), or 0.63 m SLE with MICI (95th percentile; [[#DeConto--2021|DeConto et al., 2021]] ). ISMIP6 ensemble ranges are wider for the high scenarios (RCP8.5/SSP5-8.5) than the low (RCP2.6/SSP1-2.6), in part because more simulations were available. The ISMIP6 simulations that apply an ice-shelf collapse scenario based on exceedance of a surface meltwater threshold ( [[#Trusel--2015|Trusel et al., 2015]] ), driven by CMIP5 models, show only a small increase in mass loss (around 0–0.04 m), mostly from the Peninsula, due in part to the small number of ice shelves predicted to collapse this century ( [[#Seroussi--2020|Seroussi et al., 2020]] ). Simulations driven by the CMIP5 model HadGEM2-ES, which has unusually extreme warming in the Ross Sea ( [[#Barthel--2020|Barthel et al., 2020]] ), show a larger mass loss (up to about 0.05 m) in East Antarctica under ice-shelf collapse ( [[#Edwards--2021|Edwards et al., 2021]] ). The ISMIP6 projections do not include the efficient meltwater drainage or atmospheric feedbacks that could reduce mass loss further ( [[#Seroussi--2020|Seroussi et al., 2020]] ).

The relationship between emissions scenario and AIS response varies across the studies, with emulated ISMIP6 projections showing a slight negative scenario dependence in the median (–0.01 m) from SSP1-2.6 to SSP5-8.5, and LARMIP-2-based projections showing a slight positive scenario-dependence in the median (0.02 m; Table 9.3). A lack of clear scenario dependence in the median masks large individual variations across climate and ice-sheet models, whereby the net AIS contribution response to emissions scenario depends on the relative magnitudes of the atmosphere, ocean and ice-sheet responses ( [[#Barthel--2020|Barthel et al., 2020]] ; [[#Seroussi--2020|Seroussi et al., 2020]] ; [[#Edwards--2021|Edwards et al., 2021]] ). Climate and ice-sheet models do not project that the AIS response will be the same under high or low greenhouse gas emissions in 2100; rather, there is no consensus on the sign of the change. In contrast, strong scenario dependence is seen from RCP4.5 to RCP8.5 in projections that allow MICI ( [[#9.4.2.4|Section 9.4.2.4]] ; [[#DeConto--2021|DeConto et al., 2021]] ), though less so than earlier projections ( [[#DeConto--2016|DeConto and Pollard, 2016]] ) due to later ice-shelf disintegrations. A negative or positive scenario dependence of the AIS response this century cannot be deduced from recent observations, because there is still ''low confidence'' in attributing the causes of observed mass loss ( [[#9.4.2.1|Section 9.4.2.1]] ), and neither regional mass increases by SMB nor regional mass losses by ice flow have a linear relationship with global mean temperature (Sections 9.4.2.1, 9.4.2.2, 9.4.2.3). There is therefore ''low agreement'' on the relationship between emissions scenario and AIS response. However, in the longer term, mass loss is expected to dominate ( [[#9.4.2.6|Section 9.4.2.6]] ).

The LARMIP-2 median projections are higher than those of the ISMIP6 emulator (by 0.04–0.07 m), and the 95th percentiles are two to three times higher. Two possible reasons for the differences between the emulated ISMIP6 and LARMIP-2 projections are assessed: the set of ice-sheet models (Annex II) and the parameter values determining sub-shelf melt sensitivity to ocean temperature ( [[#9.4.2.3|Section 9.4.2.3]] ; Box 9.3). Using only the 13 ice-sheet models common to ISMIP6 and LARMIP-2 reduces the LARMIP-2 median projections by 0.02–0.03 m SLE and the 95th percentiles by 0.04–0.08 m SLE (Table 9.3). This approximately halves the difference in medians, but has a relatively small effect on the upper end. Sub-shelf melt sensitivity has a larger effect, due to the wide variation of estimates from different regions and methods. Using only the Pine Island Glacier sub-shelf melt distribution (Sections 9.4.2.2 and 9.4.2.3) in the ISMIP6 emulator gives a median Antarctic projection of about 0.08 m in 2100 in all scenarios before historical correction, compared with around 0 m using only the mean Antarctic distribution; the published projections use a joint distribution ( [[#Edwards--2021|Edwards et al., 2021]] ). [[#Reese--2020|Reese et al. (2020)]] find that using the basal melt sensitivities of LARMIP-2 yields an order of magnitude greater mass loss under RCP8.5 than with the ISMIP6 mean Antarctic values. Halving the basal melt sensitivity parameter range (i.e., in line with a continental mean estimate: [[#9.4.2.3|Section 9.4.2.3]] ) would lead to a halving of the LARMIP-2 dynamic contribution. This would reconcile the LARMIP-2 and ISMIP6 emulator median and 95th percentile projections using the common subset of models within about 0.02–0.05 m. There is therefore ''limited evidence'' that the ISMIP6 and LARMIP-2 projections could be reconciled by using common ice-sheet models and basal melt sensitivity values.

It is not possible to distinguish which of ISMIP6 and LARMIP-2 is more realistic, due to limitations in historical simulations (Box 9.3) and understanding of basal melting ( [[#9.4.2.3.2|Section 9.4.2.3.2]] ), so the projections are combined using a ‘p-box’ approach ( [[#9.6.3.2|Section 9.6.3.2]] ). The mean of the ISMIP6 emulated and LARMIP-2 medians gives the assessed median projections, and the outer edges of the 17–83% ranges give the outer edges of the assessed ''likely'' (17–83%) ranges – that is, encompassing the structural and parametric uncertainties of both methods, giving ''medium confidence'' in their combined projections. The main difference between this assessment and SROCC is to increase the medians of the lower scenarios by 0.05–0.07 m, so that all SSPs are similar to SROCC assessment of RCP8.5, and to substantially increase the upper ends of the ''likely'' ranges: by 0.14–0.16 m for RCP2.6/SSP1-2.6 and RCP4.5/SSP2-4.5, and 0.06 m for RCP8.5/SSP5-8.5. The increase relative to SROCC is partly due to the increase in LARMIP-2 projections relative to the original LARMIP study ( [[#Levermann--2014|Levermann et al., 2014]] ), arising from the larger number of participating ice-sheet models ( [[#Levermann--2020|Levermann et al., 2020]] ). The historical dynamic response to pre-2015 climate forcing applied to the ISMIP6 emulator could be overestimated, due to the assumption of a constant future rate (Box 9.3). This assessment encompasses SROCC and all projections since, except the 83rd percentiles of projections that allow MICI under RCP8.5 ( [[#DeConto--2021|DeConto et al., 2021]] ) and the Structured Expert Judgement (SEJ) under 5°C shown in SROCC ( [[#Bamber--2019|Bamber et al., 2019]] ). Both are used in further p-box estimates to give the outer limits of ''low'' ''confidence'' assessments ( [[#9.6.3.2|Section 9.6.3.2]] ).

In summary, it is ''likely'' that the AIS will continue to lose mass throughout this century under all emissions scenarios – that is, dynamic losses driven by ocean warming and ice-shelf disintegration will ''likely'' continue to outpace increasing snowfall ( ''medium confidence'' ). The upper end of projections is not well constrained, due to different assumptions about the future sensitivity of sub-shelf basal melting to ocean warming and the proposed marine ice cliff instability triggered by ice-shelf disintegration (Sections 9.4.2.3 and 9.4.2.4; Box 9.4).

<div id="_idContainer049" class="Basic-Text-Frame"></div>

'''Table 9.3''' '''|''' '''Projected sea level contributions in metres from the Antarctic Ice Sheet in 2100 relative to 199''' '''5–2''' '''014, unless otherwise stated, for selected Representative Concentration Pathway (RCP) and Shared Socio-economic Pathways (SSP) scenarios.''' Italics denote partial contributions. The historical dynamic response omitted from ISMIP6 simulations is estimated to be 0.33 ± 0.16 mm yr <sup>–1</sup> (0.03 m ± 0.01 m in 2100 relative to 2015; Box 9.3). The climate forcing is described in Supplementary Material 7.SM.2.

{| class="wikitable"
|-
| colspan="5"| '''Representative Concentration Pathways (RCPs)'''

|-
| '''Study'''

| '''RCP2.6'''

| '''RCP4.5'''

| '''RCP8.5'''

| '''Notes'''

|-
| IPCC AR5 ( [[#Church--2013b|Church et al., 2013b]] )

| 0.06 (–0.04 to +0.16)

| 0.05 (–0.05 to +0.15)

| 0.04 (–0.08 to +0.14)

| Median and ''likely'' (≥ 66% range) contribution

|-
| IPCC SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] )

| 0.04 (0.01 to 0.11)

| 0.06 (0.01 to 0.15)

| 0.12 (0.03 to 0.28)

| Median and ''likely'' (66% range) contribution. Combination of five studies

|-
| ''ISMIP6 CMIP5-forced'' ( [[#Seroussi--2020|Seroussi et al., 2020]] ) ''; excludes historical dynamic response''

| ''–0.01 to +0.16''

| ''–''

| ''–0.08 to +0.30''

| ''Range of ISMIP6 multi-model contributions in 2100 relative to 2015 from 2 ESMs for RCP2.6 and 6 ESMs for RCP8.5''

|-
| ''LARMIP-2; excludes surface mass balance (SMB)'' ( [[#Levermann--2020|Levermann et al., 2020]] )

| ''0.13 (0.07 to 0.24)''

''[0.04 to 0.37]''

| ''0.14 (0.07 to 0.28)''

''[0.05 to 0.44]''

| ''0.17 (0.09 to 0.36)''

''[0.06 to 0.58]''

| ''Median (67% range) [90% range] LARMIP-2 multi-model dynamic contribution in 2100 relative to 1900''

|-
| MICI ( [[#DeConto--2021|DeConto et al., 2021]] )

| 0.08 (0.06 to 0.12)

[0.06 to 0.15]

| 0.09 (0.07 to 0.11)

[0.07 to 0.15]

| 0.34 (0.19 to 0.53)

[0.11 to 0.63]

| Median (66% range) [90% range]

|-
| colspan="5"|
|-
| colspan="5"| '''Shared Socio-economic Pathways (SSPs)'''

|-
| '''Study'''

| '''SSP1-2.6'''

| '''SSP2-4.5'''

| '''SSP5-8.5'''

| '''Notes'''

|-
| colspan="5"| Multi-model ensemble projections

|-
| ''ISMIP6 CMIP6-forced'' ( [[#Payne--2021|Payne et al., 2021]] ) ''; excludes historical dynamic response''

| ''–0.05 to +0.01''

| ''–''

| ''–0.09 to +0.11''

| ''Range of ISMIP6 multi-model contributions in 2100 relative to 2015 from 1 ESM for SSP1-2.6 and 4 ESMs for SSP5-8.5''

|-
| ISMIP6 all (CMIP5 and CMIP6-forced) including historical dynamic response

| –0.05 (0.04 to 0.08)

[0.03 to 0.11]

| ''–''

| 0.04 (0.00 to 0.12) [–0.02 to +0.23]

| Median (66% range) [90% range] contribution from ISMIP6 CMIP5 and CMIP5-forced multi-model ensembles, (see caption)

|-
| ''Emulated ISMIP6; excludes historical dynamic response'' ( [[#Edwards--2021|Edwards et al., 2021]] )

| ''0.04 (–0.01 to +0.10)''

''[–0.05 to +0.14]''

| ''0.04 (–0.02 to +0.10)''

''[–0.06 to +0.14]''

| ''0.04 (–0.01 to +0.09)''

''[–0.05 to +0.14]''

| ''Median (66% range) [90% range] contribution in 2100 relative to 2015 from emulator of ISMIP6 used with [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] climate forcing''

|-
| '''Emulated ISMIP6 total'''

| '''0.09 (0.03 to 0.14)'''

'''[–0.01 to +0.19]'''

| '''0.09 (0.03 to 0.14)'''

'''[–0.01 to +0.18]'''

| '''0.08 (0.03 to 0.14)'''

'''[0.00 to 0.18]'''

| '''Emulated ISMIP6, but relative to 1995–2014 and including historical dynamic response (see caption)'''

|-
| ''SMB''

| ''–0.02 (–0.03 to –0.01)''

''[–0.04 to –0.01]''

| ''–0.03 (–0.04 to –0.02)''

''[–0.06 to –0.01]''

| ''–0.05 (–0.07 to –0.03)''

''[–0.09 to –0.02]''

| ''Median (66% range) [90% range] SMB estimated for the AR5, used to correct LARMIP-2 below''

|-
| ''LARMIP-2; excludes SMB''

| ''0.15 (0.08 to 0.29)''

''[0.05 to 0.44]''

| ''0.17 (0.09 to 0.33)''

''[0.06 to 0.49]''

| ''0.20 (0.10 to 0.39)''

''[0.07 to 0.61]''

| ''Median (66% range) [90% range] dynamic contribution from LARMIP-2 multi-model method used with [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] climate forcing''

|-
| ''LARMIP-2 subset of models; excludes SMB''

| ''0.14 (0.08 to 0.26) [0.05 to 0.39]''

| ''0.15 (0.08 to 0.29) [0.05 to 0.45]''

| ''0.17 (0.10 to 0.35) [0.06 to 0.54]''

| ''As above, but using only the 13 of 16 ice-sheet models common to both ISMIP6 and LARMIP-2''

|-
| ''LARMIP-2 subset of models; includes SMB''

| ''0.11 (0.05 to 0.24) [0.03 to 0.37]''

| ''0.12 (0.05 to 0.26) [0.02 to 0.42]''

| ''0.12 (0.05 to 0.30) [0.01 to 0.49]''

| ''As above, but including the SMB estimate''

|-
| '''LARMIP-2 total'''

| '''0.13 (0.06 to 0.27)'''

'''[0.03 to 0.41]'''

| '''0.14 (0.06 to 0.29)'''

'''[0.02 to 0.46]'''

| '''0.15 (0.05 to 0.34)'''

'''[0.01 to 0.57]'''

| ''Median (66% range) [90% range] dynamic contribution from LARMIP-2 multi-model method used with [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] climate forcing, including the SMB estimate''

|-
| '''This assessment: combination of emulated ISMIP6 and LARMIP-2'''

| '''0.11 (0.03 to 0.27)'''

'''[–0.01 to +0.41]'''

| '''0.11 (0.03 to 0.29)'''

'''[–0.01 to +0.46]'''

| '''0.12 (0.03 to 0.34)'''

'''[0.00 to 0.57]'''

| '''Median (66% range) [90% range] assessment combining emulated ISMIP6 and LARMIP-2'''

|}

<div id="9.4.2.6" class="h3-container"></div>

<span id="projections-beyond-2100-1"></span>
==== 9.4.2.6 Projections Beyond 2100 ====

<div id="h3-28-siblings" class="h3-siblings"></div>

The SROCC assessed the median and ''likely'' range of Antarctic SLE contributions at 2300 as 0.16 (0.07–0.37) m under RCP2.6 and 1.46 (0.60–2.89) m under RCP8.5, based on three studies. It was noted that ''deep uncertainty'' remained beyond 2100: while solid Earth feedbacks could reduce ice loss over multi-century time scales, MICI ( [[#9.4.2.4|Section 9.4.2.4]] ) might give contributions higher than the ''likely'' ranges. The SROCC also presented structured expert judgement (SEJ) projections for comparison ( [[#Bamber--2019|Bamber et al., 2019]] ), which give higher values. Since SROCC, three studies have made projections to 2300: (i) [[#Rodehacke--2020|Rodehacke et al. (2020)]] assessed two methods for implementing precipitation changes (based on repeating 2071–2100 forcings beyond 2100), which both gave negative projections at 2300 because the dynamic response was very small (–0.11 to –0.01 m SLE for RCP2.6; –0.25 to –0.07 m for RCP8.5 forcing); (ii) In contrast, simulations forced by 2081–2100 ocean-only projections under RCP8.5/SSP5-8.5 beyond 2100, using two implementations of the ISMIP6 ‘non-local’ basal melt parametrizations (Box 9.3 and [[#9.4.2.2|Section 9.4.2.2]] ) and two sliding laws, are all positive (0.08 m to 0.96 m SLE by 2300), though these do not include the negative contribution from SMB changes ( [[#Lipscomb--2021|Lipscomb et al., 2021]] ); (iii) Finally, [[#DeConto--2021|DeConto et al. (2021)]] update projections for the MICI hypothesis ( [[#9.4.2.4|Section 9.4.2.4]] ) using the extensions of the RCPs to 2300, and obtain far higher contributions: median (17–83%) ranges of 1.09 (0.71–1.35) m SLE under RCP2.6 and 9.60 (6.87–13.54) m SLE under RCP8.5. These are larger than previous estimates ( [[#DeConto--2016|DeConto and Pollard, 2016]] ), particularly at the upper end: 0.68 (0.29–1.13) m SLE for RCP2.6 and 8.40 (7.47–9.76) m for RCP8.5 ( [[#Edwards--2019|Edwards et al., 2019]] ), which can largely be explained by the higher maximum ice cliff calving rate. LARMIP-2 dynamic projections (Box 9.3) are also estimated under the extended SSPs and corrected with SMB (as in [[#9.4.2.5|Section 9.4.2.5]] ), giving median (17–83%) ranges of 0.40 (0.18–0.78) m SLE at 2300 under SSP1-2.6 and 1.57 (0.68–3.14) m under SSP5-8.5. The longer time scale may invalidate the linear response assumption of LARMIP-2, which neglects any self-dampening or self-amplifying processes. The ranges of projections for 2300 without MICI ( [[#Golledge--2015|Golledge et al., 2015]] ; [[#Bulthuis--2019|Bulthuis et al., 2019]] ; [[#Levermann--2020|Levermann et al., 2020]] ; [[#Rodehacke--2020|Rodehacke et al., 2020]] ; [[#Lipscomb--2021|Lipscomb et al., 2021]] ; ‘assessed ice-sheet contributions’ in [[#9.6.3.5|Section 9.6.3.5]] are –0.14 to +0.78 m SLE under RCP2.6/SSP1-2.6, and –0.27 to 3.14 m SLE under RCP8.5/SSP5-8.5). The lower bounds are the 5th percentile of [[#Bulthuis--2019|Bulthuis et al. (2019)]] and the lowest mean/median from [[#Rodehacke--2020|Rodehacke et al. (2020)]] , respectively; the upper bounds are the 83% percentiles of the LARMIP-2 estimates. These ranges are wider than SROCC ''likely'' ranges, and more consistent with the SEJ ( [[#Bamber--2019|Bamber et al., 2019]] ). However, projections in which Antarctica contributes much more than the assessed ranges under sustained very high greenhouse gas emissions – that is, around 7–14 m to GMSL by 2300 ( [[#DeConto--2021|DeConto et al., 2021]] ), cannot be ruled out, and are taken as a sensitivity case ( [[#9.6.3.5|Section 9.6.3.5]] ; Table 9.11). In summary, there is ''high confidence'' that Antarctic mass loss will be greater beyond 2100 under high greenhouse gas emissions, but the large range of projections mean we have only ''low confidence'' in the likely AIS contribution to GMSL by 2300 for a given scenario. ''Deep uncertainty'' remains in the role of AIS instabilities under very high emissions.

The West and East Antarctic ice sheets are considered to be tipping elements – that is, susceptible to critical thresholds. The SR1.5 ( [[#Hoegh-Guldberg--2018|Hoegh-Guldberg et al., 2018]] ) assessed that a threshold for WAIS instability may be close to 1.5°C–2°C ( ''medium confidence'' ), as only RCP2.6 led to long-term projections of less than 1 m ( [[#Golledge--2015|Golledge et al., 2015]] ; [[#DeConto--2016|DeConto and Pollard, 2016]] ). Based on the agreement of a further study ( [[#Bulthuis--2019|Bulthuis et al., 2019]] ), SROCC confirmed that low emissions would limit Antarctic ice loss over multi-century time scales ( ''high confidence'' ), but it was not possible to determine whether this was sufficient to prevent substantial ice loss ( ''medium confidence'' ). Since SROCC, new studies have revisited this topic ( [[#Garbe--2020|Garbe et al., 2020]] ; [[#Rodehacke--2020|Rodehacke et al., 2020]] ; [[#Van%20Breedam--2020|Van Breedam et al., 2020]] ; [[#DeConto--2021|DeConto et al., 2021]] ; [[#Lipscomb--2021|Lipscomb et al., 2021]] ), allowing a more complete assessment along with other studies ( [[#Feldmann--2015|Feldmann and Levermann, 2015]] ; [[#Clark--2016|Clark et al., 2016]] ; [[#Golledge--2017a|Golledge et al., 2017a]] ; [[#Edwards--2019|Edwards et al., 2019]] ) and the extension to LARMIP-2 above. The majority project 0–1.3 m SLE on multi-century time scales under scenarios of 1°C–2°C warming. Projections can increase up to 2 m SLE under high basal melt sensitivity to ocean warming ( [[#9.4.2.3|Section 9.4.2.3]] ; [[#Lipscomb--2021|Lipscomb et al., 2021]] ) or MICI ( [[#9.4.2.4|Section 9.4.2.4]] ). On multi-millennial time scales (≥2,000 years), many projections remain below 1.6 m SLE under 1°C–2°C warming – that is, less than about half of the WAIS in SLE (see also [[#9.6.3.5|Section 9.6.3.5]] and Figure 9.30). Other studies project majority or total loss of WAIS under 1°C–2°C warming, exceeding 2 m SLE, under the higher end of the warming range (≥1.5°C), or high ocean warming (≥0.5°C) and/or high basal melting around WAIS, or MICI. All but two of these multi-millennial studies use variants of the same ice-sheet model, though different modelling choices mean they can be considered quasi-independent. Simulations of previous interglacial periods often show near or total WAIS disintegration, with mass loss exceeding 3 m SLE (e.g. Figure 9.18), although limitations of these studies or inferences that can be drawn under different forcings limit confidence in the robustness of these as quantitative analogues (Sections 9.4.2.4 and 9.6.2). Overall, increased evidence and agreement on the time scales and drivers of mass loss confirm the SR1.5 assessment that a threshold for WAIS instability may be close to 1.5°C–2°C ( ''medium confidence'' ), and that the probability of passing a threshold is larger for 2°C warming than for 1.5°C ( ''medium confidence'' ), particularly under strong ocean warming. New projections agree with previous studies that only part of WAIS would be lost on multi-century time scales if warming remains less than 2°C ( ''medium confidence'' ). There is ''limited agreement'' about whether complete disintegration would eventually occur at this level of warming, but ''medium confidence'' this would take millennia.

Under around 2°C–3°C peak warming, complete or near-complete loss of the WAIS is projected in most studies after multiple millennia ( ''low confidence'' ), with continent-wide mass losses of around 2–5 m SLE or more; this could occur on multi-century time scales under very high basal melting ( [[#Lipscomb--2021|Lipscomb et al., 2021]] ) or widespread ice-shelf loss and/or MICI ( ''low confidence'' ) ( [[#Sun--2020|Sun et al., 2020]] ; [[#DeConto--2021|DeConto et al., 2021]] ). Mass losses under around 2°C–3°C warming could be less than 2 m SLE, particularly for multi-century time scales, low basal melting, or less responsive sliding laws. If warming exceeds around 3°C above pre-industrial, part of the EAIS (typically the Wilkes Subglacial Basin) is projected to be lost on multi-millennial time scales ( ''low confidence'' ), with total AIS mass loss equivalent to around 6–12 m or more sea level rise; mass loss could be much smaller if the dynamic response is small ( [[#Bulthuis--2019|Bulthuis et al., 2019]] ; [[#Rodehacke--2020|Rodehacke et al., 2020]] ), or much faster under widespread ice-shelf loss and/or MICI ( [[#Sun--2020|Sun et al., 2020]] ; [[#DeConto--2021|DeConto et al., 2021]] ). A study by [[#Garbe--2020|Garbe et al. (2020)]] suggests that 6°C sustained warming and associated mass loss of about 12 m SLE may be a critical threshold beyond which the ice sheet reorganizes to a new state, leading to large losses from East Antarctica (including the Aurora Subglacial Basin) and leading to a further 10 m sea level contribution per degree of warming; other studies also show much higher mass loss per °C at higher levels of warming ( [[#9.6.3.5|Section 9.6.3.5]] and Figure 9.30; [[#Van%20Breedam--2020|Van Breedam et al., 2020]] ; [[#DeConto--2021|DeConto et al., 2021]] ).

The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ; [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ) assessed that Antarctic mass losses could be irreversible over decades to millennia ( ''low confidence'' ). [[#Garbe--2020|Garbe et al. (2020)]] show that the AIS is always volumetrically smaller when regrowing under a given warming level than when it retreats under the same forcing. Even if retreat followed by regrowth results in a net zero change in volume, the spatial distribution of mass may be altered, especially in parts of West Antarctica vulnerable to MISI. Projections that start reducing CO <sub>2</sub> concentrations from 2030 onwards, reaching pre-industrial levels around 2300, show sea level contributions exceeding 1 m by 2500 when including MICI ( [[#DeConto--2021|DeConto et al., 2021]] ). New research therefore confirms SROCC assessment that mass loss from the AIS is irreversible on decadal to millennial time scales ( ''low confidence'' ) (FAQ 9.1), and suggests that reducing atmospheric CO <sub>2</sub> concentrations or temperatures to pre-industrial levels may not be sufficient to prevent or reverse substantial Antarctic mass losses ( ''low confidence'' ).

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