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

==== 9.6.3.5 Multi-century and Multi-millennial Sea Level Rise ====

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Neither AR5 nor SROCC discussed the sea level commitment associated with historical emissions. Since AR5, new evidence has suggested that historical emissions up to 2016 will lead to a ''likely'' committed sea level rise (i.e., the rise that would occur in the absence of additional emissions) of 0.7–1.1 m up to 2300, while pledged emissions through 2030 increase the committed rise to 0.8–1.4 m ( [[#Nauels--2019|Nauels et al., 2019]] ).

Between the baseline period (1995–2014) and 2300, AR5 projected a GMSL rise of 0.38–0.82 m under a non-specific low-emissions scenario and 0.9–3.6 m under a non-specific high-emissions scenario (Table 9.11). The SROCC projected 0.6–1.0 m under RCP2.6 and 2.3–5.3 m under RCP8.5 ( ''low confidence'' ). RCP-based projections for 2300 published since AR5 span a broader range, even excluding studies employing SEJ or MICI, with 17th–83rd percentile projections ranging from 0.3–2.9 m for RCP2.6 and 1.7–6.8 m for RCP8.5 (Table 9.SM.8; [[#Kopp--2014|Kopp et al., 2014]] , 2017; [[#Nauels--2017|Nauels et al., 2017]] , 2019; [[#Bamber--2019|Bamber et al., 2019]] ; [[#Palmer--2020|Palmer et al., 2020]] ). Conservatively extending the ISMIP6- and LARMIP-2-based projections beyond 2100 by assuming no subsequent change in ice-sheet mass flux rates (an approach similar to that adopted by [[#Palmer--2020|Palmer et al. (2020)]] for the Greenland Ice Sheet and for the Antarctic Ice Sheet dynamics) leads to a GMSL change up to 2300 of 0.8–2.0 m under SSP1-2.6 and 1.9–4.1 m under SSP5-8.5 (17th–83rd percentile), while incorporating the ice-sheet contributions for 2300 assessed in [[#9.4.1.4|Section 9.4.1.4]] and [[#9.4.2.6|Section 9.4.2.6]] leads to 0.6–1.5 m and 2.2–5.9 m, respectively. Incorporating Antarctic results from a model with MICI ( [[#9.4.2.4|Section 9.4.2.4]] ), using RCP forcing to inform SSP-based projections, leads to 1.4–2.1 m for SSP1-2.6 and 9.5–16.2 m for SSP5-8.5 ( [[#DeConto--2021|DeConto et al., 2021]] ). Incorporating the SEJ-based ice-sheet projections of [[#Bamber--2019|Bamber et al. (2019)]] for 2°C and 5°C stabilization scenarios yields 1.0–3.1 m for SSP1-2.6, and 2.4–6.3 m for SSP5-8.5, although because of the differences in scenarios, the SSP1-2.6 estimates may be overestimated and the SSP5-8.5 may be underestimated. The eightfold uncertainty range across projection methods under SSP5-8.5 reflects ''deep uncertainty'' in the multi-century response of ice sheets to strong climate forcing.

Taking into account all these approaches, including published projections for RCP2.6, under SSP1-2.6 GMSL will rise between 0.3 and 3.1 m by 2300 ( ''low confidence'' ). This projection range indicates that, while SROCC projections under low emissions to 2300 are consistent with no ice-sheet acceleration after 2100, there is the possibility of a much broader range of outcomes at the high end, reflected in the range of published GMSL projections. Under SSP5-8.5, GMSL will rise between 1.7 and 6.8 m by 2300 in the absence of MICI and by up to 16 m considering MICI, a wider range than AR5 or SROCC assessments, but consistent with published projections ( ''low confidence'' ).

On still longer time scales, AR5 concluded with ''low confidence'' that the multi-millennial GMSL commitment sensitivity to warming was about 1–3 m °C <sup>–1</sup> GSAT increase. Two process-model studies since AR5 ( [[#Clark--2016|Clark et al., 2016]] ; [[#Van%20Breedam--2020|Van Breedam et al., 2020]] ) indicate higher commitments (Figure 9.30). Ice sheets dominate the multi-millennial sea level commitment (Sections 9.4.1.4 and 9.4.2.6), but the two studies disagree on the relative contribution of the Greenland and Antarctic ice sheets. Notably, processes such as MICI ( [[#9.4.2.4|Section 9.4.2.4]] ) that are a major factor behind the ''deep uncertainty'' in century-scale AIS response do not appear to have a substantial effect on the multi-millennial magnitude ( [[#DeConto--2016|DeConto and Pollard, 2016]] ). Only one of the studies of multimillennial GMSL commitments includes scenarios consistent with 1.5°C of peak warming ( [[#Clark--2016|Clark et al., 2016]] ); this study suggests a 2000-year commitment at 1.5°C of about 2.3–3.1 m, with approximately an additional 1.4–2.3 m commitment between 1.5°C and 2.0°C (i.e., about 3 to 5 m °C <sup>–1</sup> ). Taken together, both studies show a 2000-year GMSL commitment of about 2–6 m for peak warming of about 2°C, 4–10 m for 3°C, 12–16 m for 4°C, and 19–22 m for 5°C ( ''medium agreement'' , ''limited evidence'' ) (Table 9.10). GMSL rise continues after 2000 years, leading to a 10,000-year commitment of about 6–7 m for 1.5°C of peak warming (based on [[#Clark--2016|Clark et al., 2016]] ), and based on both studies of about 8–13 m for 2.0°C, 10–24 m for 3.0°C, 19–33 m for 4.0°C, and 28–37 m for 5°C ( ''medium agreement'' , ''limited evidence'' ) (Table 9.10).

An indicative metric for the equilibrium sea level response can be provided by comparing paleo GSAT and GMSL during past multimillennial warm periods (Sections 2.3.1.1, 2.3.3.3 and 9.6.2; Figure 9.9). However, caution is needed as the present and past warm periods differ in astronomical and other forcings (Cross-chapter Box 2.1) and in terms of polar amplification. The Last Interglacial ( ''likely'' 5–10 m higher GMSL than today and 0.5°C–1.5°C warmer than 1850–1900; [[#9.6.2|Section 9.6.2]] ; Table 9.6) is consistent with the [[#Clark--2016|Clark et al. (2016)]] projections for the 10,000-year commitment associated with 1.5°C of warming. Similarly, the Mid-Pliocene Warm Period ( ''very likely'' 5–25 m higher GMSL than today and ''very likely'' 2.5°C–4°C warmer) ( [[#9.6.2|Section 9.6.2]] ; Table 9.6) is consistent with the range of 10,000-year commitments associated with 2.5–4°C of warming, but GMSL reconstructions provide only a weak, broad constraint on model-based projections. An additional paleo constraint comes from the Early Eocene Climatic Optimum, which indicates that 10–18°C of warming is associated with ice-free conditions and a ''likely'' GMSL rise of 70–76 m (Sections 2.3.3 and 9.6.2). Together with model-based projections ( [[#Clark--2016|Clark et al., 2016]] ; [[#Van%20Breedam--2020|Van Breedam et al., 2020]] ), this period suggests that commitment to ice-free conditions would occur for peak warming of about 7°C–13°C ( ''medium agreement,'' ''limited evidence'' ).

On the basis of modelling studies, paleo constraints, single-ice-sheet studies finding multimillennial nonlinear responses from both the Greenland and Antarctic ice sheets (Sections 9.4.1.4 and 9.4.2.6), and the underlying physics, we conclude that GMSL commitment is nonlinear in peak warming on time scales of both 2,000 and 10,000 years ( ''medium confidence)'' and exceeds the AR5 assessment of 1–3 m °C <sup>–1</sup> ( ''medium agreement'' , ''limited evidence'' ) (Table 9.9). Although thermosteric sea level will start to decline slowly about 2,000 years after emissions cease, the slower responses from the Greenland and Antarctic ice sheets mean that GMSL will continue to rise for 10,000 years under most scenarios ( ''medium confidence'' ).

Since AR5, a small number of modelling studies have examined the reversibility of the multimillennial sea level commitment under carbon dioxide (CO <sub>2</sub> ) removal, solar radiation modification or local ice shelf engineering. The slow response of the deep ocean to forcing leads to global-mean thermosteric sea level fall occurring long afterward, even if CO <sub>2</sub> levels are restored after a transient increase: global mean thermosteric sea level rise takes more than a millennium to reverse ( [[#Ehlert--2018|Ehlert and Zickfeld, 2018]] ). Rapid reversion to pre-industrial CO <sub>2</sub> concentrations has been found to be ineffective at fostering regrowth of the AIS ( [[#DeConto--2021|DeConto et al., 2021]] ) but may reduce the multimillennial sea level commitment ( [[#DeConto--2016|DeConto and Pollard, 2016]] ). Altering sub-ice-shelf bathymetry ( [[#Wolovick--2018|Wolovick and Moore, 2018]] ) or triggering ice shelf advance through massive snow deposition ( [[#Feldmann--2019|Feldmann et al., 2019]] ) might interrupt marine ice sheet instability ( [[#9.4.2.4|Section 9.4.2.4]] ) and thus reduce sea level commitment. A reversion to pre-industrial Greenland Ice Sheet temperatures with solar radiation modification is projected to stop mass loss in Greenland but leads to minimal regrowth ( [[#Applegate--2015|Applegate and Keller, 2015]] ). Based on ''limited evidence'' , carbon dioxide removal, solar radiation modification, and local ice-shelf engineering may be effective at reducing the yet-to-be-realized sea level commitment, but ineffective at reversing GMSL rise ( ''low confidence'' ).

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'''Table 9.11''' '''|''' '''Global mean sea level (GMSL) projections between 199''' '''5–2''' '''014 and 2300 for total change and individual contributions. Low emissions projections from: AR5 ( [[#Church--2013b|Church et al., 2013b]] ); RCP2.6 from SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ) and published projections (Table 9.SM.8); and SSP1-2.6 (from this Report). High emissions projections from: AR5 ( [[#Church--2013b|Church et al., 2013b]] ); RCP8.5 from SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ) and published projections (Table 9.SM.8); and SSP5-8.5 (this Report).''' Values for AR5 ( [[#Church--2013b|Church et al., 2013b]] ) and SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ) are adjusted from the 1986–2005 baseline used in past reports. Only total values are shown for published ranges. Only the Antarctic contribution changed between AR5 ( [[#Church--2013b|Church et al., 2013b]] ) and SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ). If a range is given, it is the 17th–83rd percentile range.

{| class="wikitable"
|-
|
| '''Low'''

| colspan="2"| '''RCP2.6'''

| colspan="4"| '''SSP1-2.6'''

|-
| m relative to 1995–2014

| '''AR5'''

| '''SROCC'''

| '''Post-AR5 Published Range'''

| '''No Ice-sheet Acceleration After 2100'''

| '''Assessed Ice-sheet Contribution'''

| '''MICI'''

| '''SEJ'''

|-
| '''Thermal expansion'''

| colspan="2"| 0.07–0.46 m

|
| colspan="4"| 0.19–0.35 m

|-
| '''Greenland'''

| colspan="2"| 0.14 m

|
| 0.22–0.39 m

| colspan="2"| 0.11–0.25 m

| 0.28–1.28 m

|-
| '''Antarctica'''

| colspan="2"| 0.21–0.25 m

|
| –0.05 to +1.14 m

| –0.14 to +0.78 m

| 0.71–1.35 m

| –0.11 to +1.56 m

|-
| '''Glaciers'''

| colspan="2"| n/a

|
| colspan="4"| 0.12–0.29 m

|-
| '''Land-water storage'''

| –0.03 m

| 0.07–0.37 m

|
| colspan="4"| 0.05–0.10 m

|-
|
|-
| '''Total (2300)'''

| 0.38–0.82 m

| 0.57–1.04 m

| 0. 3–2.9 m

| 0.8–2.0 m

| 0.6–1.5 m

| 1.4–2.1 m

| 1. 0–3.1 m

|-
|
|-
|
| '''High'''

| colspan="2"| '''RCP8.5'''

| colspan="4"| '''SSP5-8.5'''

|-
| m relative to 1995–2014

| '''AR5'''

| '''SROCC'''

| '''Post-AR5 Published Range Without (with) MICI'''

| '''No Ice-Sheet Acceleration after 2100'''

| '''Assessed Ice-sheet Contribution'''

| '''MICI'''

| '''SEJ'''

|-
| '''Thermal expansion'''

| colspan="2"| 0.28–1.80 m

|
| colspan="4"| 0.92–1.51 m

|-
| '''Greenland'''

| colspan="2"| 0.30–1.18 m

|
| 0.53–0.88 m

| colspan="2"| 0.32–1.75 m

| 0.40–2.23 m

|-
| '''Antarctica'''

| 0.02–0.19 m

| 0.60–2.89 m

|
| –0.39 to +1.55 m

| –0.28 to +3.13 m

| 6.87–13.54 m

| 0.03–3.05 m

|-
| '''Glaciers'''

| colspan="2"| 0.29–0.39 m

|
| colspan="4"| 0.32 m

|-
| '''Land-water storage'''

| colspan="2"| n/a

|
| colspan="4"| 0.05–0.10 m

|-
|
|-
| '''Total (2300)'''

| 0.89–3.56 m

| 2.25– 5.34 m

| 1.7–6.8 (up to 14.1) m

| 1.7–4.0 m

| 2.2–5.9 m

| 9.5–16.2 m

| 2.4– 6.3 m

|}

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[[File:844eaaa3c81cd3a84b1f38cbbbb50487 IPCC_AR6_WGI_Figure_9_30.png]]

'''Figure 9.30''' '''|''' '''Global mean sea level (GMSL) commitment as a function of peak global surface air temperature.''' From models ( [[#Clark--2016|Clark et al., 2016]] ; [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#Garbe--2020|Garbe et al., 2020]] ; [[#Van%20Breedam--2020|Van Breedam et al., 2020]] ) and paleo data on 2000-year '''(lower row)''' and 10,000 year '''(upper row)''' time scales. Columns indicate different contributors to GMSL rise (from left to right: total GMSL change, Antarctic Ice Sheet, Greenland Ice Sheet, global mean thermosteric sea level rise, and glaciers). Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

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'''Box 9.4 | High-end Storyline of 21st-century Sea Level Rise'''

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In this box, we outline a storyline (Glossary, Box 10.2; [[#Shepherd--2018|Shepherd et al., 2018]] ) for high-end sea level projections for 2100. This storyline considers processes whose quantification is highly uncertain regarding the timing of their possible onset and/or their potential to accelerate sea level rise. These processes are therefore not considered for the assessed upper bound of ''likely'' sea level rise by 2100 in section 9.6.3.3, as the ''likely'' range includes only processes that can be projected skilfully with at least ''medium confidence'' (based on ''agreement'' and ''evidence'' ).

As noted by SROCC, stakeholders with a low risk tolerance (e.g., those planning for coastal safety in cities and long-term investment in critical infrastructure) may wish to consider global-mean sea level rise above the assessed ''likely'' range by the year 2100, because ‘ ''likely'' ’ implies an assessed likelihood of up to 16% that sea level rise by 2100 will be higher (see also [[#Siegert--2020|Siegert et al., 2020]] ). Because of our limited understanding of the rate at which some of the governing processes contribute to long-term sea level rise, we cannot currently robustly quantify the likelihood with which they can cause higher sea level rise before 2100 ( [[#Stammer--2019|Stammer et al., 2019]] ).

In light of such ''deep uncertainty'' , we employ a storyline approach in examining the potential for, and early warning signals of a high-end sea level scenario unfolding within this century. In doing so, we note upfront that the main uncertainty related to high-end sea level rise is ‘when’ rather than ‘if’ it arises: the upper limit of 1.01 m of ''likely'' sea level range by 2100 for the SSP5-8.5 scenario will be exceeded in any future warming scenario on time scales of centuries to millennia ( ''high confidence'' ), but it is uncertain how quickly the long-term committed sea level will be reached ( [[#9.6.3.5|Section 9.6.3.5]] ). Hence, global mean sea level might rise well above the ''likely'' range before 2100, which is reflected by assessments of ice-sheet contributions based on structured expert judgement ( [[#Bamber--2019|Bamber et al., 2019]] ) leading to a 95th percentile of projected future sea level rise as high as 2.3 m in 2100 ( [[#9.6.3.3|Section 9.6.3.3]] ).

A plausible storyline for such high-end sea level rise in 2100 assumes a strong warming scenario ( [[IPCC:Wg1:Chapter:Chapter-4#4.8|Section 4.8]] ). The storyline considers faster-than-projected disintegration of marine ice shelves and the abrupt, widespread onset of marine ice cliff instability (MICI) and marine ice sheet instability (MISI) in Antarctica ( [[#9.4.2.4|Section 9.4.2.4]] ), and faster-than-projected changes in both the surface mass balance and dynamical ice loss in Greenland. While conceptual studies provide ''medium evidence'' of these processes, substantial uncertainties and ''low agreement'' in quantifying their future evolution arise from limited process understanding, limited availability of evaluation data, missing or crude representation in model simulations, their high sensitivity to uncertain boundary conditions and parameters, and/or uncertain atmosphere and ocean forcing (Sections 9.4.1.2; 9.4.2.2).

In Antarctica, high warming might lead to floating ice shelves starting to break up earlier than expected due to processes not yet accounted for in ice-sheet models or in current climate models used to force ice-sheet projections. Such processes include hydrofracturing driven by surface meltwater, and increase in ocean thermal forcing driven by ocean circulation changes (Sections 9.2.2.3, 9.2.3.2 and 9.4.2.3; [[#Hellmer--2012|Hellmer et al., 2012]] , 2017; [[#Silvano--2018|Silvano et al., 2018]] ; [[#Hazel--2020|Hazel and Stewart, 2020]] ). In particular, the Thwaites and Pine Island Glacier ice shelves could potentially disintegrate this century, which might trigger MICI before 2100 ( [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#DeConto--2021|DeConto et al., 2021]] ). MISI could potentially develop earlier and faster than simulated by the majority of models if fast flowing ice streams follow plastic, instead of currently assumed more viscous, sliding laws ( [[#Sun--2020|Sun et al., 2020]] ). Oceanic feedbacks could drive high-end sea level rise by changes in the meltwater-driven overturning circulation in ice cavities that cause additional melting ( [[#Jeong--2020|Jeong et al., 2020]] ); by a warming of the ocean water in contact with the ice shelves due to increased stratification and thus reduced vertical mixing (Sections 9.2.2.3 and 9.2.3.2; [[#Golledge--2019|Golledge et al., 2019]] ; [[#Moorman--2020|Moorman et al., 2020]] ; [[#Sadai--2020|Sadai et al., 2020]] ); or by an increase in sea ice cover due to increased ocean stratification ( [[#9.3.2.1|Section 9.3.2.1]] ), which could reduce the amount of warm, moist air that reaches the continent, and limit the mass gain from snowfall over the ice sheet ( [[#Sadai--2020|Sadai et al., 2020]] ).

In Greenland, stronger mass loss than currently projected might also occur ( [[#Aschwanden--2019|Aschwanden et al., 2019]] ; [[#Khan--2020|Khan et al., 2020]] ; T. [[#Slater--2020|]] [[#Slater--2020|Slater et al., 2020]] ). For example, warming-induced dynamical changes in atmospheric circulation could enhance summer blocking and produce more frequent extreme melt events over Greenland similar to the record mass loss of more than 500 Gt in summer 2019 ( [[#9.4.1.1|Section 9.4.1.1]] ; [[#Delhasse--2018|Delhasse et al., 2018]] ; [[#Sasgen--2020|Sasgen et al., 2020]] ). Cloud processes in polar areas that are not well represented in models could further enhance surface melt ( [[#Hofer--2019|Hofer et al., 2019]] ), as could feedbacks between surface melt and the increasing albedo from meltwater, detritus and pigmented algae ( [[#9.4.1.1|Section 9.4.1.1]] ; [[#Cook--2020|Cook et al., 2020]] ). The same ice dynamical processes associated with basal melt and MISI discussed for Antarctica could also occur in Greenland, as long as the ice sheet is in contact with the ocean.

The strength of all these processes is currently understood to depend strongly on global mean temperature and polar amplification, with additional linkages through feedback from global mean sea level ( [[#Gomez--2020|Gomez et al., 2020]] ). These dependencies on a joint forcing imply that processes are strongly correlated. Hence, both their uncertainties and their possible cascading contribution to high-end sea level rise are expected to combine. Therefore, high-end sea level rise can occur if one or two processes related to ice-sheet collapse in Antarctica result in an additional sea level rise at the maximum of their plausible ranges (Sections 9.4.2.5 and 9.6.3.3; Table 9.7) or if several of the processes described in this box result in individual contributions to additional sea level rise at moderate levels. In both cases, global-mean sea level rise by 2100 would be substantially higher than the assessed ''likely'' range, as indicated by the projections including ''low confidence'' processes reaching in 2100 as high as 1.6 m at the 83rd percentile and 2.3 m at the 95th percentile ( [[#9.6.3.3|Section 9.6.3.3]] ).

Identifying the potential drivers of a high-end sea level rise allows identification of sites and observables that can provide early warnings of a much faster sea level rise than the ''likely'' range of this and previous reports. One potential site for such monitoring is Thwaites Glacier, which is melting faster in some places and slower in others than models simulate. At this glacier, the effect of tides and channelling of warm water flows on the melting is evident ( [[#Milillo--2019|Milillo et al., 2019]] ), making the floating ice shelf potentially vulnerable to breakup from hydrofracturing, driven by surface meltwater, much earlier than expected. In addition, the glacier is retreating towards a zone with deeper bedrock, which at its present rate of retreat would be reached in 30 years ( [[#Yu--2019|Yu et al., 2019]] ). Thwaites Glacier is therefore a strong candidate to experience large-scale MISI and/or MICI ( [[#Golledge--2019|Golledge et al., 2019]] ; [[#DeConto--2021|DeConto et al., 2021]] ), making it the ideal site for monitoring early warning signals of accelerated sea level rise from Antarctica. Such signals could possibly be observed within the next few decades ( [[#Scambos--2017|Scambos et al., 2017]] ).

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