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

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

<div id="h3-14-siblings" class="h3-siblings"></div>

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]] ).

<div id="_idContainer033" class="Basic-Text-Frame"></div>

[[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'' ).

<div id="9.3" class="h1-container"></div>

<span id="sea-ice-1"></span>