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

=== 9.2.4 Steric and Dynamic Sea Level Change ===

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==== 9.2.4.1 Global Mean Thermosteric Sea Level Change ====

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Changes in globally averaged ocean heat content (OHC) cause global mean thermosteric sea level (GMTSL) change (Box 9.1). The observed increased OHC for 1971–2018 of 325 to 546 ZJ ( ''very likely'' range) ( [[IPCC:Wg1:Chapter:Chapter-7#7.2|Section 7.2]] , Box 7.2) has led to a GMTSL rise of 0.03 to 0.06 m out of a total global mean sea level (GMSL) of 0.07 to 0.15 m ( ''very likely'' range) ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] , Tables 2.7 and 9.5, and Cross-Chapter Box 9.1).

Projections of GMTSL rise in AR5 ( [[#Church--2013b|Church et al., 2013b]] ) and SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ) were derived from the CMIP5 ensemble, after removing drift estimated based on pre-industrial control simulations. Differences between removing a linear and a quadratic drift are small ( [[#Hobbs--2016a|Hobbs et al., 2016a]] ; [[#Hermans--2021|Hermans et al., 2021]] ). These prior assessments filled in projections for models that did not provide GMTSL rise for all scenarios, by calculating the heat content of the climate system from global surface air temperature and net radiative flux, then converting this to GMTSL rise using each model’s diagnosed expansion efficiency coefficient. In AR5, the associated uncertainties were derived by assuming a normal distribution, with the 5th–95th percentile CMIP5 ensemble range taken as the ''likely'' range (±1 standard deviation).

In this Report, global surface air temperature projections are not derived directly from the CMIP6 ensemble (Box 4.1). Therefore, to produce projections of OHC and GMTSL rise consistent with the Report’s assessment of equilibrium climate sensitivity and transient climate response ( [[IPCC:Wg1:Chapter:Chapter-7#7.5.2.2|Section 7.5.2.2]] ), this chapter employs a two-layer energy budget emulator (Supplementary Materials 7.SM.2, 9.SM.4.3). Since AR5, climate model emulators have been increasingly used to predict GMTSL (Cross-Chapter Box 7.1; [[#Kostov--2014|Kostov et al., 2014]] ; [[#Palmer--2018|Palmer et al., 2018]] , 2020; [[#Nauels--2019|Nauels et al., 2019]] ). The expansion efficiency coefficient that relates GMTSL and OHC for the two-layer emulator has a mean and standard deviation of 0.113 ± 0.013 m YJ <sup>–1</sup> (Supplementary Material 9.SM.4.3). This approach yields a ''likely'' thermosteric contribution between 1995–2014 and 2100 that represents a minimal change from AR5 and SROCC (Table 9.8). The two-layer emulator GMTSL projected median and 17th–83rd percentile, or ''likely'' , range is 0.12 (0.09 to 0.15) m for SSP1-1.9, 0.14 (0.11 to 0.18) m for SSP1-2.6, 0.20 (0.16 to 0.24) m for SSP2-4.5, 0.25 (0.21 to 0.30) m for SSP3-7.0, and 0.30 (0.24 to 0.36) m for SSP5-8.5 by 2100 ( [[#9.6.3.2|Section 9.6.3.2]] and Tables 9.1, 9.8 and 9.9). The two-layer model heat content increases slightly faster than that of the total depth CMIP6 ensemble, which is related to its role in the assessed energy balance (Section 7.SM.2), but with a similar ensemble spread (Table 9.1). Projecting the ''likely'' factor by which 1995–2014 to 2081–2100 OHC change exceeds change over 1971 to 2018 in CMIP6 yields 3 to 5 for SSP1-2.6, 4 to 6 for SSP2-4.5, 5 to 7 for SSP3-7.0, and 5 to 8 for SSP5-8.5. The two-layer model ''likely'' equivalents are 2 to 3 for SSP1-2.6, 3 to 4 for SSP2-4.5, 4 to 5 for SSP3-7.0, and 4 to 6 for SSP5-8.5.

For reconstructions, the expansion efficiency coefficient is required for the conversion between ocean temperature and steric sea level over a specific time scale. Combining the assessed sea level and energy data over 1995 to 2014 (drawn from the analysis in Cross-Chapter Box 9.1) results in a coefficient of 0.1210 ± 0.0014 m YJ <sup>–1</sup> , or 0.6607 ± 0.0076 m °C <sup>–1</sup> in terms of mean ocean temperature. The two-layer emulator assessment used in AR6 results in 0.113 ± 0.013 m YJ <sup>–1</sup> , or 0.617 ± 0.071 m °C <sup>–1</sup> (Appendices 7.SM.2, 9.SM.4). Both of these estimates are in line with an independent estimate of 0.70 m/°C ( [[#Hieronymus--2019|Hieronymus, 2019]] ) and other estimates, for example, 0.116 ± 0.011 m YJ <sup>–1</sup> ( [[#Kuhlbrodt--2012|Kuhlbrodt and Gregory, 2012]] ), but are significantly larger than the temperature to sea level conversion used in AR5 (0.42 m °C <sup>–1</sup> based on SST and the estimated range from [[#Levermann--2013|Levermann et al., 2013]] ). The expansion coefficient is not fixed across models, nor in time, as it varies depending on which water masses are storing the added heat, and the commitment time scale ( [[#Hallberg--2013|Hallberg et al., 2013]] ). For paleoclimate, a scaling for sea surface temperature (0.6 m °C <sup>–1</sup> ) or global surface air temperature (GSAT; see Cross-Chapter Box 2.3) can be estimated, but mean ocean temperature is in phase with steric sea level change, while sea surface temperatures are not (Figure 9.9; [[#Shakun--2012|Shakun et al., 2012]] ; [[#Tierney--2020|Tierney et al., 2020]] ). Thus, while conversions between OHC, mean ocean temperature and GMTSL across applications are within uncertainty ranges ( ''medium confidence'' ) (Table 9.1), little consistency is found when correlating these variables to SST or GSAT, which may vary independently.

Short-lived climate forcers (Sections 6.3 and 6.6.3) are associated with a sea level commitment, due to an OHC and mean ocean temperature response that lasts substantially longer than their atmospheric forcing and SST response, although not as long as the sea level commitment associated with CO <sub>2</sub> emissions (Sections 9.2.1.1 and 4.4.4). For example, [[#Zickfeld--2017|Zickfeld et al. (2017)]] find that about 70% of the thermosteric sea level rise associated with methane forcing would persist 100 years after the elimination of methane emissions, and 40% would persist for more than 500 years.

In summary, consistent relationships between OHC ( [[#9.2.2.1|Section 9.2.2.1]] ), mean ocean temperature and GMTSL are found using two-layer emulators, CMIP6 models, and modern and paleo observations to provide ''medium confidence'' in the 0.113 ± 0.013 m YJ <sup>–1</sup> , or 0.617 ± 0.071 m °C <sup>–1</sup> ''likely'' ranges of assessed conversion values. It is possible to estimate relationships between SST or GSAT change and GMTSL rise, but conversions are not generally applicable and depend on time scale and application.

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'''Table 9.1''' '''|''' '''Projected contributions to median and 1''' '''7–8''' '''3% (parentheses) and''' '''5–9''' '''5% [square brackets] ranges of thermosteric sea level from AR5 ( [[#Church--2013b|Church et al., 2013b]] ), CMIP6 ( [[#Jevrejeva--2020|Jevrejeva et al., 2020]] ; [[#Hermans--2021|Hermans et al., 2021]] ) and the two-layer energy balance model (described in Sections 7.SM.2, 9.SM.4 and Box 4.1) averaged over 208''' '''1–2''' '''100, with respect to a baseline of 199''' '''5–2''' '''014.''' Note that AR5 and SROCC interpret 5–95% range as the ''likely'' range, while in this table square brackets are used for consistency.

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

| '''RCP2.6/SSP1-2.6'''

| '''RCP4.5/SSP2-4.5'''

| '''RCP8.5/SSP5-8.5'''

|-
| '''IPCC AR5 and SROCC GMTSL'''

'''( [[#Church--2013b|Church et al., 2013b]] ; [[#Oppenheimer--2019|Oppenheimer et al., 2019]]''' )

| 0.13 [0.09 to 0.17] m

| 0.18 [0.13 to 0.22] m

| 0.26 [0.20 to 0.32] m

|-
| '''CMIP6 5–95% GMTSL'''

'''( [[#Hermans--2021|Hermans et al., 2021]] )'''

| 0.14 [0.08 to 0.17] m

| 0.18 [0.11 to 0.23] m

| 0.26 [0.17 to 0.33] m

|-
| '''CMIP6 5–95% GMTSL'''

'''( [[#Jevrejeva--2020|Jevrejeva et al., 2020]] )'''

| –

| 0.19 [0.13 to 0.24] m

| 0.27 [0.19 to 0.35] m

|-
| '''Assessed GMTSL based on two-layer model 17–83% and 5–95% (Sections''' '''7.SM.2''' ''', 9.SM.4)'''

| 0.13 (0.11 to 0.16) [0.09 to 0.19] m

| 0.17 (0.14 to 0.21)

[0.12 to 0.25] m

| 0.25 (0.20 to 0.30)

[0.18 to 0.35] m

|-
| '''Total OHC 17–83% and 5–95% from assessed two-layer model (Sections''' '''7.SM.2''' ''', 9.SM.4)'''

| 1.18 (0.99 to 1.42)

[0.86 to 1.65] YJ

| 1.56 (1.33 to 1.86)

[1.19 to 2.12] YJ

| 2.23 (1.92 to 2.64)

[1.71 to 3.00] YJ

|-
| '''0–2000 m OHC 17–83% and 5–95% from CMIP6 (Figure 9.6)'''

| 1.06 (0.80 to 1.31)

[0.66 to 1.64] YJ

| 1.35 (1.08 to 1.67)

[0.90 to 1.84] YJ

| 1.89 (1.60 to 2.29)

[1.28 to 2.58] YJ

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==== 9.2.4.2 Ocean Dynamic Sea Level Change ====

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Projections of ocean dynamic sea level change (Box 9.1) on multi-annual time scales resemble the patterns of steric sea level change in the open ocean (Figures 9.11 and 9.12; [[#Lowe--2006|Lowe and Gregory, 2006]] ; [[#Pardaens--2011|Pardaens et al., 2011]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). On shorter time scales, especially in extratropical coastal areas, there may be an important barotropic component (also called bottom pressure change) due mostly to changes in wind-driven circulation and eddies apparent in the variance of ocean dynamic sea level (Figure 9.12; [[#Roberts--2016|Roberts et al., 2016]] ; [[#Hughes--2018|Hughes et al., 2018]] ). This component is highly sensitive to ocean model resolution ( [[#Chassignet--2020|Chassignet et al., 2020]] ). Steric sea level change is associated with local changes in temperature and salinity, which come about through changes in surface fluxes of heat and freshwater ( [[#9.2.1.2|Section 9.2.1.2]] ) and through redistribution of existing water masses by changed ocean circulation and mixing processes (Figure 9.12 and Sections 9.2.2.1 and 9.2.3). Redistribution of water masses often involves anticorrelated thermosteric and halosteric changes (Figure 9.12), especially in the Atlantic ( [[#Pardaens--2011|Pardaens et al., 2011]] ; [[#Bouttes--2014|Bouttes et al., 2014]] ; [[#Durack--2014|Durack et al., 2014]] ; [[#Griffies--2014|Griffies et al., 2014]] ; [[#Han--2017|Han et al., 2017]] ).

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[[File:4df7eb594a13c19fc82467942cad5f6b IPCC_AR6_WGI_Figure_9_12.png]]

'''Figure 9.12''' '''|''' '''(a–f) Coupled Model Intercomparison Project Phase 6 (CMIP6) multi-model mean projected change contributions to relative sea level change in (a, d) steric sea level anomaly, (b, e) thermosteric sea level anomaly, and (c, f) halosteric sea level anomaly between 199''' '''5–2''' '''014 and 208''' '''1–2''' '''100 using a method that does not require a reference level ( [[#Landerer--2007|Landerer et al., 2007]] ).''' Global mean change has been removed from these figures, consistent with the methods in Sections 9.6.3 and 9.SM.4 and the definitions of [[#Gregory--2019|Gregory et al. (2019)]] . ( [[#Gregory--2019|Gregory et al., 2019]] ). See Figure 9.27 for global mean sea level (GMSL). (g–i) Standard deviation of ocean dynamic sea level change from (g) Aviso observations (10-day high-pass filter); (h) five-day mean of high-resolution Ocean Model Intercomparison Project phase 2 (OMIP-2) models forced with observed fluxes; and (i) five-day mean of low-resolution OMIP-2 models which are comparable in resolution to the models in (a–f). No overlay indicates regions with high model agreement, where ≥80% of models agree on the sign of change. Diagonal lines indicate regions with low model agreement, where &lt;80% of models agree on the sign of change (see Cross-Chapter Box Atlas.1 for more information). Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

Ocean dynamic sea level change is strongly affected by internal variability ( [[#9.6.1.4|Section 9.6.1.4]] ), partly from interannual to decadal coupled atmosphere–ocean modes of variability via wind-driven redistribution (Annex IV; [[#Griffies--2014|Griffies et al., 2014]] ; [[#Han--2017|Han et al., 2017]] ) and partly from intrinsic ocean variability, particularly in higher-resolution simulations (such as HighResMIP), which statistically resemble observations, even on short time scales (Figure 9.12; [[#Griffies--2014|Griffies et al., 2014]] ; [[#Sérazin--2016|Sérazin et al., 2016]] ; [[#Llovel--2018|Llovel et al., 2018]] ; [[#Chassignet--2020|Chassignet et al., 2020]] ). High-resolution simulations are not used in relative sea level projections ( [[#9.6.3|Section 9.6.3]] ) due to the limited range of forcing scenarios. The most marked feature of long-term regional sea level change in the continuous satellite altimetry record, beginning in 1992, is the east–west dipole in the Pacific Ocean (rising more rapidly in the east, see also [[#9.6.1.3|Section 9.6.1.3]] ), which persisted until 2015, and can be explained by anomalously strong trade winds ( [[#Merrifield--2012|Merrifield et al., 2012]] ; [[#England--2014|England et al., 2014]] ; [[#Griffies--2014|Griffies et al., 2014]] ; [[#Takahashi--2016|Takahashi and Watanabe, 2016]] ; [[#Han--2017|Han et al., 2017]] ) together with associated changes in surface heat flux ( [[#Piecuch--2019|Piecuch et al., 2019]] ). The most notable features of sub-annual variability in altimetry are eddies and tides, which are directly simulated only in high-resolution models ( [[#Haigh--2019|Haigh et al., 2019]] ; [[#Chassignet--2020|Chassignet et al., 2020]] ).

Projections of the pattern and amplitude of regional ocean dynamic sea level change in CMIP6 and previous model generations show a large model spread, of a similar size to the geographical spread (Figure 9.12). The model spread derives from model dependence of changes both in surface fluxes ( [[#9.2.1.2|Section 9.2.1.2]] ) and in the ocean response ( [[#9.2.2|Section 9.2.2]] ). The spread is similar in CMIP6 and CMIP5, and is largest in regions with large projected variations in ensemble-mean ocean dynamic sea level change ( [[#Lyu--2020a|Lyu et al., 2020a]] ), such as the Southern Ocean Dipole with an ocean dynamic sea level rise north of the ACC and a fall to the south, the Atlantic Dipole with a sea level rise north of 40°N and a fall in 20°N–40°N, the Northwest Pacific Dipole, and the large sea level rise in the Arctic ( [[#Church--2013b|Church et al., 2013b]] ; [[#Slangen--2014a|Slangen et al., 2014a]] , 2015; [[#Bilbao--2015|Bilbao et al., 2015]] ; [[#Gregory--2016|Gregory et al., 2016]] ; [[#Chen--2019|]] [[#Chen--2019|C. Chen et al., 2019]] ; [[#Lyu--2020a|Lyu et al., 2020a]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). Patterns of change are consistent between model simulations and observations ( ''medium confidence'' ). The major model ensemble-mean features resemble thermosteric sea level change, as expected from altered input of heat to the ocean without changing circulation, while model spread results from the diversity in redistribution of the heat content of the unperturbed ocean ( [[#9.2.2.1|Section 9.2.2.1]] ; [[#Bouttes--2014|Bouttes and Gregory, 2014]] ; [[#Gregory--2016|Gregory et al., 2016]] ; [[#Huber--2017|Huber and Zanna, 2017]] ; [[#Lyu--2020b|Lyu et al., 2020b]] ; [[#Todd--2020|Todd et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ).

The Southern Ocean Meridional Dipole is driven by a northward advection of excess heat (from changes in surface fluxes) by the wind-driven circulation followed by subduction or diffusive uptake in mid-latitudes, northward redistribution of existing heat by the strengthening of that circulation, and the meridional contrast in thermal expansivity due to its temperature-dependence ( [[#Armour--2016|Armour et al., 2016]] ; [[#Gregory--2016|Gregory et al., 2016]] ; [[#Lyu--2020b|Lyu et al., 2020b]] ; [[#Todd--2020|Todd et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ).

The positive Arctic ocean dynamic sea level change is driven by increased freshwater input ( [[#Couldrey--2021|Couldrey et al., 2021]] ). The Northwest Pacific Dipole is driven by the intensification of the Kuroshio Current in response to reduced heat loss and in some models to wind stress change ( [[#Chen--2019|]] [[#Chen--2019|C. Chen et al., 2019]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ).

The North Atlantic sea level change dipole is forced by a reduction in heat loss from the ocean north of 40°N (i.e., net heat uptake), which in all Earth system models leads to a weakening of the AMOC, although the magnitude has a large model spread ( [[#9.2.3.1|Section 9.2.3.1]] ; [[#Gregory--2016|Gregory et al., 2016]] ; [[#Huber--2017|Huber and Zanna, 2017]] ). The reduced northward transport of warm, salty water ( [[#9.2.2|Section 9.2.2]] ) causes further ocean dynamic sea level change, whose details are model-dependent. North of 40°N, this redistribution leads to a sea level rise, predominantly halosteric, reinforcing the thermosteric effect of heat uptake ( [[#Couldrey--2021|Couldrey et al., 2021]] ). Comparison of observed Atlantic OHC for 1955–2017 with a reconstruction assuming no change in circulation indicates that the thermosteric sea level change resulting from southward redistribution of heat may be detectable ( [[#Zanna--2019|Zanna et al., 2019]] ). This redistribution causes a tendency for SST cooling north of 40°N and anomalous heat input from the atmosphere, and thus a positive feedback on AMOC weakening ( [[#Winton--2013|Winton et al., 2013]] ; [[#Gregory--2016|Gregory et al., 2016]] ; [[#Todd--2020|Todd et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). Many climate and ocean models agree that the AMOC weakening is associated with pronounced thermosteric sea level rise along the American coast around 40°N (Figures 9.12 and 9.26), leading to a relatively large ocean dynamic sea level rise in this region ( [[#Yin--2012|Yin, 2012]] ; [[#Bouttes--2014|Bouttes et al., 2014]] ; [[#Slangen--2014b|Slangen et al., 2014b]] ; [[#Little--2019|Little et al., 2019]] ; [[#Lyu--2020a|Lyu et al., 2020a]] ).

In summary, ocean dynamic sea level change involves changes to temperature and salinity and responses of currents to changing forcing, with significant variability driven by unforced oceanic variability. Projections of dynamic sea level variability require fully three-dimensional ocean models, and only high-resolution ocean models are statistically consistent on short time scales with satellite altimeter observations ( ''very high confidence'' ).

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