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

== 9.2 Oceans ==

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=== 9.2.1 Ocean Surface ===

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==== 9.2.1.1 Sea Surface Temperature ====

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The IPCC Fifth Assessment Report (AR5; [[#Hartmann--2013|Hartmann et al., 2013]] ) assessed that it is ''virtually certain'' that global sea surface temperature (SST) has increased since the beginning of the 20th century ( ''very high confidence'' ). The Special Report on Ocean and Cryosphere in a Changing Climate (SROCC) did not assess past SST change. Since AR5, improvements in the understanding of recent SST biases in the observational records, especially extending ship-based observations with buoy-based observations and improved treatment of sea ice, have had important consequences for key climate change indicators such as global mean surface temperature (GMST), global surface air temperature (GSAT), and SST (Cross-Chapter Box 2.3). The AR5 assessment is confirmed, and it is now ''very likely'' that global mean SST changed by 0.88 [0.68 to 1.01] °C from 1850–1900 to 2011–2020, and 0.60 [0.44 to 0.74] °C from 1980 to 2020 (Figure 9.3 and Table 2.4).

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[[File:c664066722b424728e3f3413047341a5 IPCC_AR6_WGI_Figure_9_3.png]]

'''Figure''' '''9.3 |''' '''Sea surface temperature (SST) and its changes with time. (a)''' Time series of global mean SST anomaly relative to 1950–1980 climatology. Shown are paleoclimate reconstructions and PMIP models, observational reanalyses (HadISST) and multi-model means from the Coupled Model Intercomparison Project (CMIP) historical simulations, CMIP projections, and HighResMIP experiment. '''(b)''' Map of observed SST (1995–2014 climatology HadISST). '''(c)''' Historical SST changes from observations. '''(d)''' CMIP 2005–2100 SST change rate. '''(e)''' Bias of CMIP. '''(f)''' CMIP change rate. '''(g)''' 2005–2050 change rate for SSP5-8.5 for the CMIP ensemble. '''(h)''' Bias of HighResMIP (bottom left) over 1995–2014. '''(i)''' HighResMIP change rate for 1950–2014. ( '''j)''' 2005–2050 change rate for SSP5-8.5 for the HighResMIP ensemble. No overlay indicates regions with high model agreement, where ≥80% of models agree on sign of change. Diagonal lines indicate regions with low model agreement, where &lt;80% of models agree on 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).

Regions vary in the rate of SST warming, with slight cooling in some regions (Figure 9.3). The SROCC ( [[#Collins--2019|Collins et al., 2019]] ) and [[IPCC:Wg1:Chapter:Chapter-7#7.4.4|Section 7.4.4]] assess SST changes over specific regions, which are consistent with the changes reported here. The tropical ocean has been warming faster than other regions since 1950, with the fastest warming in regions of the tropical Indian and western Pacific oceans (Figure 9.3), due to a combination of local atmosphere–ocean coupling, the Indonesian Throughflow ( [[#9.2.3.4|Section 9.2.3.4]] and Figure 9.11), and trends in the Walker circulation (Sections 2.3.1.4.1 and 3.3.3.1, and Figure 3.16). The western boundary currents of the subtropical gyres have warmed faster than the global mean over the past century. There remains ''low agreement'' in the changes of the location and the dynamical changes in western boundary current extensions (Sections 2.3.3.4.2 and 9.2.3.4, and Figure 9.3). In the Arctic, the mean SST increase over the last two decades is similar to, or only slightly higher than, the global average (J.-L [[#Chen--2019|]] [[#Chen--2019|Chen et al., 2019]] ). In contrast, the eastern Pacific Ocean, subpolar North Atlantic Ocean and Southern Ocean have warmed more slowly than the global average or cooled (Figure 9.3). Surface warming in the subpolar Southern Ocean has been slower than the global average since the 1950s, and this pattern is consistent with the upwelling around Antarctica renewing surface water with pre-industrial, deeper water masses ( [[#9.2.3.2|Section 9.2.3.2]] ; [[#Frölicher--2015|Frölicher et al., 2015]] ; J. [[#Marshall--2015|]] [[#Marshall--2015|Marshall et al., 2015]] ; [[#Armour--2016|Armour et al., 2016]] ). New evidence since SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) confirms slight cooling since the 1980s around the subpolar Southern Ocean, contrasting with marked warming directly northward of it ( [[#9.2.3.2|Section 9.2.3.2]] ; [[#Haumann--2020|Haumann et al., 2020]] ; [[#Rye--2020|Rye et al., 2020]] ; [[#Auger--2021|Auger et al., 2021]] ). In eastern boundary upwelling systems, SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) reported ''low agreement'' between SST trends in recent decades, due to varying spatio-temporal resolution and interannual to multi-decadal variability. Satellite evidence not included in SROCC shows that 92% of these regions warmed more slowly than neighbouring offshore locations between 1982 and 2015, so upwelling may buffer the near shore from warming ( [[#9.2.3.5|Section 9.2.3.5]] ; [[#Varela--2018|Varela et al., 2018]] ). Coupled ocean-atmospheric modes of variability strongly affect regional SST (Cross-Chapter Box 3.1 and Annex IV). In summary, a positive SST trend since 1950 is evident globally, but there is ''very high confidence'' that the Indian Ocean, western equatorial Pacific Ocean, and western boundary currents have warmed faster than the global average, while the Southern Ocean, the eastern equatorial Pacific, and the North Atlantic Ocean have warmed more slowly, or have slightly cooled.

In AR5 ( [[#Flato--2013|Flato et al., 2013]] ), a marginal improvement was noted in Coupled Model Intercomparison Project Phase 5 (CMIP5) climate model SST biases compared to Phase 3 (CMIP3) models in AR4, with a reduction in the magnitude of biases. The AR5 noted that, in several regions, large SST biases are symptomatic of errors in the representation of important processes, such as dynamics in the equatorial Pacific and North Atlantic, and Southern Ocean. Common regional biases in SST or historical SST trends are not exclusively linked to the representation of the ocean ( ''high confidence'' ), but can have multiple causes, including: errors in the representation of long-term historical trends in equatorial winds ( [[#9.2.1.2|Section 9.2.1.2]] ); misrepresentation of the forced equatorial ocean response ( [[#Karnauskas--2012|Karnauskas et al., 2012]] ; [[#Kohyama--2017|Kohyama et al., 2017]] ; [[#Coats--2018|Coats and Karnauskas, 2018]] ); thermocline depth errors ( [[#Linz--2014|Linz et al., 2014]] ); errors in atmospheric model cloud-related shortwave radiation ( [[#Hyder--2018|Hyder et al., 2018]] ); biases in ocean circulation variability ( [[#Wang--2014|]] [[#Wang--2014|C. Wang et al., 2014]] ); and deficiencies in upper ocean (Q. [[#Li--2019|]] [[#Li--2019|]] [[#Li--2019|Li et al., 2019]] ) and atmospheric ( [[#Bates--2012|Bates et al., 2012]] ) boundary layer parametrizations. In CMIP6, the mid-latitude biases in the Northern Hemisphere are improved in the multi-model mean, and the inter-model standard deviation of the zonal mean SST error is significantly decreased in the northern Hemisphere south of 50°N compared to CMIP5, though biases in equatorial regions remain essentially unchanged ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.1.1|Section 3.5.1.1]] and Figures 3.23, 3.24 and 9.3). Some long-standing ocean model biases have been reduced through increases in model resolution in CMIP6 ( [[#Bock--2020|Bock et al., 2020]] ) and improved parametrizations ( [[#Fox-Kemper--2011|Fox-Kemper et al., 2011]] ; Q. [[#Li--2016|]] [[#Li--2016|Li et al., 2016]] ; [[#Qiao--2016|Qiao et al., 2016]] ; [[#Reichl--2018|Reichl and Hallberg, 2018]] ). The High Resolution Model Intercomparison Project (HighResMIP) ensemble (Figure 9.3) has smaller cold biases in the North Atlantic and the tropical Pacific, and smaller warm biases in the upwelling regions off the western coasts of Africa, North and South America ( [[#Roberts--2018|Roberts et al., 2018]] , 2019; [[#Caldwell--2019|Caldwell et al., 2019]] ; [[#Docquier--2019|Docquier et al., 2019]] ). In summary, CMIP6 models show persistent regional biases in representing the climatological SST state ( ''very high confidence'' ), but higher resolution reduces some biases, particularly in the North Atlantic and eastern boundary upwelling systems (Figure 9.3; ''high confidence'' ).

The CMIP6 models represent the observed trends in SST patterns with greater fidelity than CMIP5, with the ocean area that is inconsistent with the observed trends decreasing by about three quarters from CMIP5 to CMIP6 ( [[#Olonscheck--2020|Olonscheck et al., 2020]] ). In some regions, the direction of SST changes in observations are consistent with CMIP6 only when including internal variability ( [[#Olonscheck--2020|Olonscheck et al., 2020]] ). This is notably the case in the equatorial Pacific, North Atlantic, and Southern Ocean, which are regions where SST is of known importance in controlling heat uptake ( [[#9.2.2.1|Section 9.2.2.1]] ) and the global radiative feedback parameter ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4.3|Section 7.4.4.3]] ). Overall, despite some persistent regional biases, CMIP6 coupled climate models reproduce the observed SST trends or high internal variability over the past century over a range of different multi-decadal periods (Figure 9.3; [[#Olonscheck--2020|Olonscheck et al., 2020]] ; [[#Watanabe--2021|Watanabe et al., 2021]] ), highlighting their skill to inform future large-scale SST changes at regional scale. Warming is projected at varying rates in all regions by 2050, except the North Atlantic Subpolar Region, the equatorial Pacific, and the Southern Ocean where models disagree ( ''high confidence'' ).

It is ''virtually certain'' that SST will continue to increase in the 21st century, at a rate depending on future emissions scenarios. The future global mean SST increase projected by CMIP6 models for the period 1995–2014 to 2081–2100 is 0.86 [5–95% range: 0.43–1.47] °C under SSP1-2.6, 1.51 [1.02 to 2.19] °C under SSP2-4.5, 2.19 [1.56 to 3.30] °C under SSP3-7.0, and 2.89 [2.01 to 4.07] °C under SSP5-8.5 (Figure 9.3). While under SSP1-2.6, the CMIP6 ensemble consistently projects that it is ''very likely'' at least 83% of the world ocean surface will have warmed by 2100, and under SSP5-8.5, at least 98% of the world ocean surface will have warmed. The spatial pattern of future change is consistent with observed SST change over the 20th century, though with notable regional differences (Figure 9.3). Long-term change in SST patterns is important for regional impacts but also affects radiative feedbacks, and therefore long-term change in climate sensitivity ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4.3|Section 7.4.4.3]] ). In the Southern Ocean, CMIP6 models project that SSTs will eventually consistently increase in the 21st century, at a rate dependent on future scenarios (Figure 9.3 and [[#9.2.3.2|Section 9.2.3.2]] ; [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ). Yet, there is only ''low confidence'' that this Southern Ocean warming will emerge by the end of the century ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4.1|Section 7.4.4.1]] ), due to the inconsistent historical and near-term simulations and observations over the 20th century (Figure 9.3). Furthermore, the equilibrium SST pattern from proxy records or simulated by climate models under CO <sub>2</sub> forcing stand in contrast with the cooling trends in the Southern Ocean observed over the past decades ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4.1.2|Section 7.4.4.1.2]] ). Similarly, the SST change pattern observed in the tropical Pacific Ocean will transition on centennial time scales to a mean pattern resembling the El Niño pattern ( ''medium confidence'' ) (Annex IV). However, it is difficult to delineate a climate change trend ressembling an El Niño pattern and El Niño variability ( [[#Wittenberg--2009|Wittenberg, 2009]] ; [[#Collins--2010|Collins et al., 2010]] ) without large ensembles ( [[#Kay--2015|Kay et al., 2015]] ). Several Pliocene SST reconstructions indicate enhanced warming in the centre of the eastern Pacific equatorial cold tongue upwelling region, consistent with reconstruction of enhanced subsurface warming and enhanced warming in coastal upwelling regions ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4.2.2|Section 7.4.4.2.2]] ). The North Atlantic subpolar gyre is projected to continue to warm more slowly than surrounding regions ( [[#Suo--2017|Suo et al., 2017]] ), as the Gulf Stream concurrently warms rapidly (Figure 9.3; [[#Cheng--2013|Cheng et al., 2013]] ) and the Atlantic Meridional Overturning Circulation further declines under greenhouse gas forcing, although models disagree about the rate of change (Figure 9.3 and [[#9.2.3.1|Section 9.2.3.1]] ). In summary, CMIP6 models show a future pattern of SST change comparable to historical trends with intensity depending on future emissions scenario, and some of the observed cooling trends over the 20th century will eventually transition to a warming SST on centennial time scales, in particular in the Southern Ocean ( ''high confidence'' ) and in the equatorial Pacific ( ''medium confidence'' ), while the North Atlantic subpolar gyre will continue to warm more slowly than the global average ( ''high confidence'' ).

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==== 9.2.1.2 Air–Sea Fluxes ====

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Air–sea fluxes of energy, freshwater, and momentum (wind stresses) are difficult to observe directly ( [[#Cronin--2019|Cronin et al., 2019]] ), so estimates of the global mean net air–sea heat flux are inferred from observed ocean warming ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.1|Section 2.3.3.1]] , Box 7.2, and Cross-Chapter Box 9.1). Air–sea heat fluxes resemble the warming patterns of CMIP3 ( [[#Domingues--2008|Domingues et al., 2008]] ; [[#Levitus--2012|Levitus et al., 2012]] ) and are consistent with the ensemble mean warming rate of CMIP5 ( [[#Cheng--2017|Cheng et al., 2017]] , 2019) and CMIP6 models ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.1.3|Section 3.5.1.3]] ). Regional air–sea fluxes in models remain a key driver of uncertainty ( [[#Huber--2017|Huber and Zanna, 2017]] ; [[#Tsujino--2020|Tsujino et al., 2020]] ). A substantial part of the upper 700 m energy increase is ''very likely'' attributed to anthropogenic forcing via increasing radiative forcing (Sections 3.5.1.3, 7.2 and 7.3).

The SROCC ( [[#Abram--2019|Abram et al., 2019]] ) and AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ) assessed that observations of air–sea fluxes had not yet reached the density or accuracy to directly detect trends beyond the noise. New evidence since SROCC confirms that direct heat and freshwater flux trends have not emerged yet as spatial (Figure 9.4), annual ( [[#Yu--2019|Yu, 2019]] ), and decadal ( [[#Zanna--2019|Zanna et al., 2019]] ) variability overwhelm detection. Since AR5, comprehensive comparisons ( [[#Bentamy--2017|Bentamy et al., 2017]] ; [[#Valdivieso--2017|Valdivieso et al., 2017]] ; [[#Yu--2017|Yu et al., 2017]] ) have used updated and new surface flux products to improve surface flux uncertainty estimates, and these comparisons note that implied global energy imbalances often exceed the observed ocean warming. Flux estimates using top of atmosphere observations and atmospheric fluxes from reanalysis have improved over past products ( [[#Trenberth--2018|Trenberth and Fasullo, 2018]] ) but require consistency adjustments ( [[#Trenberth--2019|Trenberth et al., 2019]] ) as the energy budget is not closed. Adjustments are needed for all flux products, and they remain less accurate than direct ocean heat content change measurements ( [[#Cheng--2017|Cheng et al., 2017]] ). Some regional changes are ''likely'' robust in both satellite observations and projections (Figure 9.4). Recent satellite-based surface flux products with improved retrieval algorithms and new satellites, for example, J-OFURO3 ( [[#Tomita--2019|Tomita et al., 2019]] ) and OAFlux-HR ( [[#Yu--2019|Yu, 2019]] ), provide a complete suite of turbulent fluxes including heat, moisture, and momentum. When combined with satellite-based surface radiation from Clouds and the Earth’s Radiant Energy System (CERES) Energy Balanced and Filled (EBAF; [[#Kato--2018|Kato et al., 2018]] ) and precipitation from Global Precipitation Climatology Project (GPCP; [[#Adler--2003|Adler et al., 2003]] ), full ocean-surface forcing is available since 1987 (Figure 9.4). These products agree with sparse buoy and ship observations within 30 W m <sup>–2</sup> ( [[#Bentamy--2017|Bentamy et al., 2017]] ; [[#Cronin--2019|Cronin et al., 2019]] ) ''.'' While patterns agree between models and satellites in net fluxes (Figure 9.4), the trend magnitudes are substantially weaker in models. The fluxes tending to warm the North Atlantic and Southern Ocean are consistent with the largest changes observed in the surface properties and water masses (Sections 9.2.1.1, 9.2.2.1 and 9.2.2.3). The observed trend toward a saltier Atlantic Ocean and a fresher Indian Ocean, as well as trends in evaporation minus precipitation (E-P) patterns in the equatorial Pacific (see also [[IPCC:Wg1:Chapter:Chapter-8#8.3.1|Section 8.3.1]] ) enhance the present mean pattern of wetting and drying. Elsewhere patterns are less clear, with only partial, large-scale agreement with the ‘wet gets wetter’ simplification (Sections 3.3.2.3, 4.4.1 and 4.5.1). In summary, globally integrated and large-scale fluxes are more reliably inferred from heat content and salinity change, while regional trends are rarely robust in observations; where they are robust, they tend to be underestimated or in disagreement in models ( ''very high confidence'' ).

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[[File:171fa62ce903f21d239a4870f36f3f2a IPCC_AR6_WGI_Figure_9_4.png]]

'''Figure''' '''9.4 |''' '''Global maps of observed mean fluxes (a, d, g), the observed trends in these fluxes (b, e, h) and the projected rate of change in these fluxes from SSP5-8.5 (c, f, i).''' Shown are the freshwater flux '''(a–c)''' , net heat flux '''(d–f)''' , and momentum flux or wind stress magnitude '''(g–i)''' , with positive numbers indicating ocean freshening, warming, and accelerating respectively. The means and observed trends are calculated between 1995–2014 (freshwater and wind stress) or 2001–2014 (heat). The SSP5-8.5 projected rates are between 1995–2100 using 20-year averages at each end of the time period. Observations show objective interpolation from Clouds and the Earth’s Radiant Energy System (CERES) Energy Balanced and Filled (EBAF) v4 ( [[#Kato--2018|Kato et al., 2018]] ), Objectively Analyzed air–sea Fluxes-High Resolution (OAFlux-HR) ( [[#Yu--2019|Yu, 2019]] ), and Global Precipitation Climatology Project (GPCP) ( [[#Adler--2003|Adler et al., 2003]] ) of fluxes and flux trends (b, e, h). Observed trends with no overlay indicate regions where the trends are significant at p = 0.34 level. Crosses indicate regions where trends are not significant. For (c, f, i) projections, 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).

There is ''low confidence'' in long-term wind stress trends in most regions, but a few locations have ''likely'' trends over the scatterometer era and in projections, as shown in Figure 9.4 ( [[#Desbiolles--2017|Desbiolles et al., 2017]] ; [[#Young--2019|Young and Ribal, 2019]] ; [[#Yu--2019|Yu, 2019]] ). The AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ) assessed with ''medium confidence'' that zonal wind stress over the Southern Ocean increased from the early 1980s to the 1990s ( ''medium confidence'' ) (Figure 9.4). Over 1995–2014, the zonal wind stress over the Southern Ocean continued to increase, westerly winds in the North Pacific and North Atlantic weakened, while the easterly equatorial Pacific winds of the Walker circulation strengthened (Figure 9.4). In historical simulations, CMIP5 models projected annular modes (Annex IV) to move poleward and strengthen in both hemispheres ( [[#Yang--2016|Yang et al., 2016]] ), while in CMIP6 models westerlies only strengthen over the Southern Ocean, with a weaker trend than recently observed (Figure 9.4 and Sections 4.5.1 and 4.5.3). In the tropical Pacific Ocean, a weakening trend in easterly winds and Walker circulation in the 20th century has been inferred based on observed sea level pressure data ( [[#Vecchi--2006|Vecchi et al., 2006]] ; [[#Vecchi--2007|Vecchi and Soden, 2007]] ) and coral proxies ( [[#Carilli--2014|Carilli et al., 2014]] ) and is projected to continue by CMIP6 models (Figure 9.4). Yet, over 1995–2014 observed winds have strengthened (Figure 9.4). The observed strengthening may have been influenced by a combination of factors ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4.2.1|Section 7.4.4.2.1]] ), but there is ''low confidence'' in the attribution of this signal to anthropogenic warming ( [[IPCC:Wg1:Chapter:Chapter-3#3.3.3.1|Section 3.3.3.1]] ) and ''medium confidence'' that it reflects internal variability ( [[IPCC:Wg1:Chapter:Chapter-8#8.3.2.3|Section 8.3.2.3]] ). Near-term projected changes over the Southern Ocean result from ozone recovery and greenhouse gases (Sections 4.3.3 and 4.4.3). Overall, there is only ''low confidence'' in observed and projected wind stress trends in most regions because trends in oceanic wind stresses during the satellite era have not emerged or are inconsistent with historical simulated changes.

Air–sea flux biases result from common causes in most models, and many are the same as during AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ). Important currents (e.g., Gulf Stream, Kuroshio, Antarctic Circum-polar Current patterns) are often found in erroneous locations in models, affecting SST and flux signatures ( [[#Bates--2012|Bates et al., 2012]] ; [[#Beadling--2020|Beadling et al., 2020]] ; J.-L.F. [[#Li--2020|]] [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ), but their locations are improved in high-resolution ocean models ( [[#Chassignet--2017|Chassignet et al., 2017]] , 2020; [[#Hewitt--2020|Hewitt et al., 2020]] ), and high-resolution coupled models reduce the mean air–sea flux biases ( [[#Delworth--2012|Delworth et al., 2012]] ; [[#Sakamoto--2012|Sakamoto et al., 2012]] ; [[#Small--2014|Small et al., 2014]] ; [[#Haarsma--2016|Haarsma et al., 2016]] ; [[#Caldwell--2019|Caldwell et al., 2019]] ; L.C [[#Jackson--2020|]] [[#Jackson--2020|Jackson et al., 2020]] ). Oceanic variability stems either from internal chaotic variability or atmospheric forcing ( [[#Hasselmann--1976|Hasselmann, 1976]] ; [[#Sérazin--2016|Sérazin et al., 2016]] , 2017). Large-scale variability in the ocean tends to follow atmospheric forcing in low-resolution models, while in high-resolution coupled models ocean variability drives atmospheric variability on small scales ( [[#Bishop--2017|Bishop et al., 2017]] ; [[#Small--2019|Small et al., 2019]] ), allowing these high-resolution models to mimic the coupling with clouds, precipitation, and atmospheric and oceanic boundary layers apparent in observations ( [[#Chelton--2010|Chelton and Xie, 2010]] ; [[#Frenger--2013|Frenger et al., 2013]] ). Even coarse-resolution models, such as the ocean and sea ice components used in CMIP6, show significant sensitivity in the mean and variability of SST and sea ice to modest changes in flux forcing ( [[#Tsujino--2020|Tsujino et al., 2020]] ). Finally, there is still considerable disagreement between different parametrizations of air–sea fluxes used in models and strong scatter in direct observations ( [[#Renault--2016|Renault et al., 2016]] ; [[#Brodeau--2017|Brodeau et al., 2017]] ). In summary, there is ''very high confidence'' that air–sea heat flux and stress biases are reduced in coupled models with high ocean resolution over coarse-resolution models, although the effect on trends remain unclear.

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==== 9.2.1.3 Upper-ocean Stratification and Surface Mixed Layers ====

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The density difference from surface to deep ocean is the upper-ocean stratification. The AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ) assessed that it is ''very likely'' that the thermal contribution to stratification over the fixed 0–200 m layer increased by about 1% per decade between 1971 and 2010 (based on linear trend consistently across reports). The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) found it ''very likely'' that density stratification increased by 0.46–0.51% per decade between 60°S and 60°N from 1970 to 2017). New published estimates based on a variety of different interpolated observations show that SROCC assessed rate is too low, even using the same data and methods ( [[#Li--2020|]] [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ). The 1960–2018 stratification increase is estimated at 1.2 ± 0.1% per decade from the IAP dataset, 1.2 ± 0.4% per decade from the Ishii product, 0.7 ± 0.5% per decade from the EN4 dataset, 0.9 ± 0.5% per decade from ORAS4, and 1.2 ±0.3% per decade from the National Centers for Environmental Information (NCEI) dataset (G. [[#Li--2020|]] [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ). The improved methodology for computing stratification change on individual profiles before gridding yields a global annual mean increase of 0–200 m stratification change of 0.8 ± 0.2% per decade between 1960 and 2018 ( [[#Yamaguchi--2019|Yamaguchi and Suga, 2019]] ) and a global summer mean increase of 0–200 m stratification change of 1.3 ± 0.3% per decade between 1970 and 2018 ( [[#Sallée--2021|Sallée et al., 2021]] ) is of a similar magnitude to the long-term trend ( [[#Yamaguchi--2019|Yamaguchi and Suga, 2019]] ; G. [[#Li--2020|]] [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ). In summary, there is ''limited evidence'' that focusing on changes over a fixed depth range might hide larger increases occurring at the seasonally and regionally variable pycnocline depth. There is also ''limited evidence'' that summer stratification change within the pycnocline has occurred at a rate of 8.9 ± 2.7% per decade from 1970 to 2018, and ''limited evidence'' of a winter pycnocline stratification increase ( [[#Cummins--2020|Cummins and Ross, 2020]] ; [[#Sallée--2021|Sallée et al., 2021]] ).

While AR5 and SROCC did not assess change in mixed-layer depth, the reported changes in stratification can modulate the surface mixed-layer depth, which is set by a balance between fluxes and dynamical mixing (winds, tides, waves, convection) acting against the background stratification and restratification processes (solar and dynamical). Despite the large stratification increase observed at a global scale, new evidence shows that summer mixed-layer depth deepened consistently over the globe at a rate of 2.9 ± 0.5% per decade from 1970 to 2018, with the largest deepening observed in the Southern Ocean, corresponding to overall deepening from 3–15 m per decade depending on region ( [[#Somavilla--2017|Somavilla et al., 2017]] ; [[#Sallée--2021|Sallée et al., 2021]] ). While the shorter observational record in winter (compared to summer) does not allow global winter mixed-layer trends to be reliably assessed ( [[#Sallée--2021|Sallée et al., 2021]] ), winter mixed-layer depths deepening at rates of 10 m per decade have been reported at individual long-term mid-latitude monitoring sites ( [[#Somavilla--2017|Somavilla et al., 2017]] ). Projections agree that shoaling of mixed-layer depth is expected in the 21st century, but only for strong emissions scenarios, and only in some regions (Figure 9.5). In summary, there is ''limited'' observational ''evidence'' that the mixed layer is globally deepening, while models show no emergence of a trend until later in the 21st century under strong emissions.

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[[File:ebaf291d5789321709688b78e3a6575a IPCC_AR6_WGI_Figure_9_5.png]]

'''Figure''' '''9.5 |''' '''Mixed-layer depth in (a–d) winter and (e–h) summer. (a, e)''' Observed climatological mean mixed-layer depth (based on density threshold) from the Argo Mixed Layer Depth Climatology ( [[#Holte--2017|Holte et al., 2017]] ) using observations for 2000–2019. '''(b, f)''' Bias between the observation-based estimate (2000–2019) and the 1995–2014 Coupled Model Intercomparison Project Phase 6 (CMIP6) climatological mean mixed-layer depth. '''(c, d, g, h)''' Projected mixed-layer depth (MLD) change from 1995–2014 to 2081–2100 under '''(c, g)''' SSP1-2.6 and '''(d, h)''' SSP5-8.5 scenarios. The '''(a–d)''' winter row shows December–January–February (DJF) in the Northern Hemisphere and June–July–August (JJA) in the Southern Hemisphere; The '''(e–h)''' summer row shows JJA in the Northern Hemisphere and DJF in the Southern Hemisphere. The mixed-layer depth is the depth where the potential density is 0.03 kg m <sup>–3</sup> denser than at 10 m. 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).

The SROCC assessed that upper-ocean stratification will continue to increase in the 21st century under increased radiative forcing ( ''high confidence'' ), due to increased surface temperature and high-latitude surface freshening ( [[#Bindoff--2019|Bindoff et al., 2019]] ). New climate model simulations concur with SROCC assessment of a future increase of the 0–200 m stratification under increased radiative forcing in all regions of the world ocean ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). In addition, CMIP6 climate models project a shallowing of the mixed-layer in summer and winter by the end of the century under increased radiative forcing (Figure 9.5; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ), with the exception of the Arctic showing deepening of the mixed layer as a result of sea ice retreat (Figure 9.5; [[#Lique--2018|Lique et al., 2018]] ). The regions of largest shallowing are associated with the deepest climatological mixed layer, in both winter and summer, particularly affecting the North Atlantic and the Southern Ocean basins (Figure 9.5). While CMIP6 models tend to project shallowing mixed layers under a warming climate, except at high latitudes (Figure 9.5; [[#Lique--2018|Lique et al., 2018]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ), a deepening in the summer mixed-layer depth by intensification of the surface winds and storms may explain inconsistency among models in many regions (Figure 9.5; [[#Young--2019|Young and Ribal, 2019]] ), although model mixed-layer biases are large in the summer in the Southern Ocean ( [[#Belcher--2012|Belcher et al., 2012]] ; [[#Sallée--2013a|Sallée et al., 2013a]] ; [[#Li--2016|Q. Li et al., 2016]] ; [[#Tsujino--2020|Tsujino et al., 2020]] ). Lack of observed ocean turbulence and climate model limitations do not allow for direct assessment of ocean surface turbulence change and limit confidence in past and future mixed-layer change. Understanding of turbulent processes, their representation in ocean and climate models, and their effect on mixed-layer biases have been an active and rapidly evolving topic of research since AR5 ( [[#Buckingham--2019|Buckingham et al., 2019]] ; Q. [[#Li--2019|]] [[#Li--2019|]] [[#Li--2019|Li et al., 2019]] ). Small-scale mixed-layer processes are not resolved in climate models ( [[#D’Asaro--2014|D’Asaro, 2014]] ; [[#Buckingham--2019|Buckingham et al., 2019]] ; [[#McWilliams--2019|McWilliams, 2019]] ) and despite significant improvements in their parametrization over the last decade ( [[#Fox-Kemper--2011|Fox-Kemper et al., 2011]] ; [[#Jochum--2013|Jochum et al., 2013]] ; [[#Li--2016|Q. Li et al., 2016]] , 2019; [[#Qiao--2016|Qiao et al., 2016]] ) and significant improvement in some models ( [[#Li--2017|Li and Fox-Kemper, 2017]] ; [[#Dunne--2020|Dunne et al., 2020]] ), biases in mixed-layer representation generally persist ( [[#Heuzé--2017|Heuzé, 2017]] ; [[#Williams--2018|Williams et al., 2018]] ; [[#Cherchi--2019|Cherchi et al., 2019]] ; [[#Golaz--2019|Golaz et al., 2019]] ; [[#Voldoire--2019|Voldoire et al., 2019]] ; [[#Yukimoto--2019|Yukimoto et al., 2019]] ; [[#Boucher--2020|Boucher et al., 2020]] ; [[#Danabasoglu--2020|Danabasoglu et al., 2020]] ; [[#Dunne--2020|Dunne et al., 2020]] ; [[#Kelley--2020|Kelley et al., 2020]] ). In summary, the representation of upper-ocean stratification and mixed layers has improved in CMIP6 compared to CMIP5. While it is ''virtually certain'' that the global mean upper ocean will continue to stratify in the 21st century, there is only ''low confidence'' in the future evolution of mixed-layer depth, which is projected to mostly shoal under high emissions, except in high-latitude regions where sea ice retreats.

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'''Box 9.2 | Marine Heatwaves'''

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Marine heatwaves (MHW) are periods of extreme high sea temperature relative to the long-term mean seasonal cycle ( [[#Hobday--2016|Hobday et al., 2016]] ). Studies since the Special Report on the Ocean and Cryosphere in a Changing Climate (SROCC; [[#Collins--2019|Collins et al., 2019]] ) confirm the assessment that MHW can lead to severe and persistent impacts on marine ecosystems – from mass mortality of benthic communities, including coral bleaching, changes in phytoplankton blooms, shifts in species composition and geographical distribution, and toxic algal blooms, to decline in fisheries catch and mariculture ( [[#Smale--2019|Smale et al., 2019]] ; [[#Cheung--2020|Cheung and Frölicher, 2020]] ; [[#Hayashida--2020|Hayashida et al., 2020]] ; [[#Piatt--2020|Piatt et al., 2020]] ). Unlike synoptic atmospheric heatwaves [[IPCC:Wg1:Chapter:Chapter-11#11.3|Section 11.3]] ), MHWs can extend for millions of square kilometres, persist for weeks to months, and occur at subsurface ( [[#Bond--2015|Bond et al., 2015]] ; [[#Schaeffer--2017|Schaeffer and Roughan, 2017]] ; [[#Perkins-Kirkpatrick--2019|Perkins-Kirkpatrick et al., 2019]] ; [[#Laufkötter--2020|Laufkötter et al., 2020]] ).

The SROCC established that MHWs have occurred in all basins over the last decades. Additional evidence documenting widespread occurrence of marine heat waves in all basins and marginal seas continues to accumulate (Y. [[#Li--2019|]] [[#Li--2019|]] [[#Li--2019|Li et al., 2019]] ; [[#Yao--2020|Yao et al., 2020]] ). The SROCC highlighted the role of large-scale climate modes of variability in amplifying or suppressing MHW occurrences, which has since been further corroborated, increasing confidence in climate modes as important drivers of MHWs ( [[#Holbrook--2019|Holbrook et al., 2019]] ; [[#Sen%20Gupta--2020|Sen Gupta et al., 2020]] ). More generally, understanding of processes leading to MHWs has increased since SROCC, including air–sea heat flux [[#9.2.1.2|Section 9.2.1.2]] ), increased horizontal heat advection, shoaling of the mixed-layer and suppressed mixing processes [[#9.2.1.3|Section 9.2.1.3]] ), reduced coastal upwelling and Ekman pumping [[#9.2.3.5|Section 9.2.3.5]] ), changes in eddy activities and planetary waves, and the re-emergence of warm subsurface anomalies ( [[#Holbrook--2020|Holbrook et al., 2020]] ; [[#Sen%20Gupta--2020|Sen Gupta et al., 2020]] ).

The SROCC reported with ''high confidence'' that MHWs – defined as days exceeding the 99th percentile in sea surface temperature (SST) from 1982 to 2016 – have ''very likely'' doubled in frequency between 1982 and 2016. Additional observation-based evidence and acquisition of longer observation time series since SROCC have confirmed and expanded on this assessment: since the 1980s MHWs have also become more intense and longer ( [[#Frölicher--2018|Frölicher and Laufkötter, 2018]] ; [[#Smale--2019|Smale et al., 2019]] ; [[#Laufkötter--2020|Laufkötter et al., 2020]] ). Satellite observations and reanalyses of SST show an increase in intensity of 0.04°C per decade from 1982 to 2016, an increase in spatial extent of 19% per decade from 1982 to 2016, and an increase in annual MHW days of 54% between the 1987–2016 period compared to 1925–1954 ( [[#Frölicher--2018|Frölicher et al., 2018]] ; [[#Oliver--2019|Oliver, 2019]] ). The SROCC assessed that 84–90% of all MHWs that occurred between 2006 and 2015 are ''very likely'' caused by anthropogenic warming. There is new evidence since SROCC that the frequency of the most impactful marine heatwaves over the last few decades has increased more than 20-fold because of anthropogenic global warming ( [[#Laufkötter--2020|Laufkötter et al., 2020]] ). In summary, there is ''high confidence'' that MHWs have increased in frequency over the 20th century, with an approximate doubling from 1982 to 2016, and ''medium confidence'' that they have become more intense and longer since the 1980s.

Consistent with SROCC, future MHWs are defined with reference to the historical climate conditions. The SROCC assessed that MHWs will ''very likely'' further increase in frequency, duration, spatial extent and intensity under future global warming in the 21st century. The CMIP6 projections allow us to confirm this assessment and quantify future change based on global mean probability ratio change (Box 9.2, Figure 1): they project MHWs will become four times (5–95% range: 2–9 times] more frequent in 2081–2100 compared to 1995–2014 under SSP1-2.6, or eight times (3–15 times) more frequent under SSP5-8.5. The SROCC highlighted that future change of MHWs will not be globally uniform, with the largest changes in the frequency of marine heatwaves being projected to occur in the western tropical Pacific and the Arctic Ocean ( ''medium confidence'' ). New evidence from the latest generation of climate models confirms and complements SROCC assessment (Box 9.2, Figure 1). Moderate increases are projected for mid-latitudes, and only small increases are projected for the Southern Ocean ( ''medium confidence'' ) ( [[#Hayashida--2020|Hayashida et al., 2020]] ). While under the SSP5-8.5 scenario, permanent MHWs (more than 360 days per year) are projected to occur in the 21st century in parts of the tropical ocean, the Arctic Ocean and around 45°S, the occurrence of such permanent MHWs can largely be avoided under the SSP1-2.6 scenario ( [[#Frölicher--2018|Frölicher et al., 2018]] ; [[#Oliver--2019|Oliver et al., 2019]] ; [[#Plecha--2020|Plecha and Soares, 2020]] ). The resolution of current climate models (CMIP5 and CMIP6) capture the broad features of MHWs, but they may have a bias towards weaker and longer MHWs in the historical period ( ''medium confidence'' ) ( [[#Frölicher--2018|Frölicher et al., 2018]] ; [[#Pilo--2019|Pilo et al., 2019]] ; [[#Plecha--2020|Plecha and Soares, 2020]] ) and greater intensification in western boundary current regions ( [[#Hayashida--2020|Hayashida et al., 2020]] ).

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Box 9.2

[[File:4de4e488565e4e63a395c016fcb1136e IPCC_AR6_WGI_Box_9_2_Figure_1.png]]

'''Box 9.2, Figur'''

'''e 1 |'''

'''Observed and simulated regional probability ratio of marine heatwaves (MHWs) for the 198''' '''5–2''' '''014 period and for the end of the 21st century under two different greenhouse gas emissions scenarios.''' The probability ratio is the proportion by which the number of MHW days per year has increased relative to pre-industrial times. An MHW is defined as a deviation beyond the daily 99th percentile (11-day window) in the deseasonalized sea surface temperature. '''(a)''' The MHW probability ratio from satellite observations (NOAA OISST V2.1; Huang et al. 2020) during 1985–2014. The mean warming pattern (difference in ERSST5 (Huang et al. 2017) sea surface temperature between the 1985–2014 and 1854–1900 periods) has been added to the satellite observations to calculate the probability ratio. '''(b–d)''' Coupled Model Intercomparison Project Phase 6 (CMIP6) simulated multi-model mean probability ratio of the '''(b)''' 1985–2014 period, and 2081–2100 period in the '''(c)''' SSP1 2.6 and '''(d)''' SSP5 8.5 scenarios. The areas with grey diagonal lines in (d) indicate permanent MHWs (&gt;360 heatwave days per year). These 14 CMIP6 models are included in the analysis: ACCESS-CM2, CESM2, CESM2-WACCM, CMCCCM2-SR5, CNRM-CM6-1, CNRM-ESM2-1, CanESM5, EC-Earth3, IPSL-CM6A-LR, MIROC6, MRI-ESM2-0, NESM3, NorESM2-LM, NorESM2-MM. Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

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=== 9.2.2 Changes in Heat and Salinity ===

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==== 9.2.2.1 Ocean Heat Content and Heat Transport ====

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Ocean warming – that is, increasing ocean heat content (OHC) – is an important aspect of energy on Earth: SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) reported that there is ''high confidence'' that ocean warming during 1971–2010 dominated the increase in the Earth’s energy inventory, which is confirmed by the Box 7.2 assessment that the ocean has stored 91% of the total energy gained from 1971 to 2018. As reported in Sections 2.3.3.1, 3.5.1.3 and 7.2.2.2, Box 7.2 and Cross-Chapter Box 9.1, confidence in the assessment of global OHC change since 1971 is strengthened compared to previous reports, and extended backward to include ''likely'' warming since 1871. Table 7.1 updates the estimates of total ocean heat gains from 1971 to 2018, 1993 to 2018 and 2006 to 2018. [[IPCC:Wg1:Chapter:Chapter-3#3.5.1.3|Section 3.5.1.3]] assesses that it is ''extremely likely'' that anthropogenic forcing was the main driver of the OHC increase over the historical period. [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.1|Section 2.3.3.1]] reports that current multi-decadal to centennial rates of OHC gain are greater than at any point since the last deglaciation ( ''medium confidence'' ).

Ocean warming is not uniform with depth. The AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ) assessed that, since 1971, ocean warming was ''virtually certain'' for the upper 700 m and ''likely'' for the 700–2000 m layer. Both AR5 and SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessed that the deep ocean below 2000 m had ''likely'' warmed since 1992, especially in the Southern Ocean. [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.1|Section 2.3.3.1]] provides an updated assessment of ocean temperature change for different depth layers, time periods and observation-based reconstructions (Table 2.7). [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.1|Section 2.3.3.1]] confirms the previous assessment that it is ''virtually certain'' that the upper ocean (0–700 m) has warmed since 1971, that ocean warming at intermediate depths (700–2000 m) is ''very likely'' since 2006, and that it is ''likely'' that ocean warming has occurred below 2000 m since 1992. [[IPCC:Wg1:Chapter:Chapter-3#3.5.1.3|Section 3.5.1.3]] assessed that it is ''extremely likely'' that human influence was the main driver of the ocean heat content increase observed since the 1970s, which extends into the deeper ocean ( ''very high confidence'' ), and shows that biases in potential temperature have a complex pattern (Figure 3.25). In the present section, we assess the regional patterns of this warming and associated processes driving regional ocean warming.

The rate of ocean warming varies regionally, with some regions having experienced slight cooling (Figure 9.6). The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessed that ocean warming in the 0–700 m depth is globally widespread, with slower than global average warming in the subpolar North Atlantic. The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) also estimated that the Southern Ocean accounted for around 75% of global ocean heat uptake during 1870–1995 and that 35–43% of the upper 2000 m global ocean warming occurred in the Southern Ocean over 1970–2017 (45–62% for 2005–2017). The SROCC noted that this interhemispheric asymmetry might (at least partially) be explained by high concentrations of aerosols in the Northern Hemisphere. Here, we confirm these assessments, bring new evidence attributing these regional trends, and discuss the role of decadal ocean circulation variability in redistributing heat, driving interhemispheric asymmetry of the recent rate of ocean warming ( [[#Rathore--2020|Rathore et al., 2020]] ; L. [[#Wang--2021|]] [[#Wang--2021|Wang et al., 2021]] ). Since SROCC, one new study shows that the subpolar North Atlantic ‘warming hole’ observed since the 1980s has emerged from internal climate variability and can be attributed to greenhouse gas emissions ( [[#Chemke--2020|Chemke et al., 2020]] ). A new analysis of a suite of climate models ( [[#Hobbs--2021|Hobbs et al., 2021]] ) confirms SROCC assessment, based on one paper ( [[#Swart--2018|Swart et al., 2018]] ), attributing the observed Southern Ocean warming to anthropogenic forcing. Given the large fraction of global ocean warming in the Southern Ocean and the sparse observations there before 2005, there is ''limited evidence'' that global OHC increase since 1971 might have been underestimated ( [[#Cheng--2014|Cheng and Zhu, 2014]] ; [[#Durack--2014|Durack et al., 2014]] ). Cross-Chapter Box 9.1 accounts for an increased error before 2005 in global OHC change. In summary, in the upper 2000 m since the 1970s, the subpolar North Atlantic has been slowly warming, and the Southern Ocean has stored a disproportionally large amount of anthropogenic heat ( ''medium confidence'' ).

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[[File:f6360086072bb669ef5529b8765822f8 IPCC_AR6_WGI_Figure_9_6.png]]

'''Figure 9.''' '''6 |''' '''Ocean heat content (OHC) and its changes with time. (a)''' Time series of global OHC anomaly relative to a 2005–2014 climatology in the upper 2000 m of the ocean. Shown are observations ( [[#Ishii--2017|Ishii et al., 2017]] ; [[#Baggenstos--2019|Baggenstos et al., 2019]] ; [[#Shackleton--2020|Shackleton et al., 2020]] ), model-observation hybrids ( [[#Cheng--2019|Cheng et al., 2019]] ; [[#Zanna--2019|Zanna et al., 2019]] ), and multi-model means from the Coupled Model Intercomparison Project Phase 6 (CMIP6) historical (29 models) and Shared Socio-economic Pathway (SSP) scenarios (label subscripts indicate number of models per SSP). '''(b–g)''' Maps of OHC across different time periods, in different layers, and from different datasets/experiments. Maps show the CMIP6 ensemble bias and observed ( [[#Ishii--2017|Ishii et al., 2017]] ) trends of OHC for '''(b, c)''' 0–700 m for the period 1971–2014, and '''(e, f)''' 0–2000 m for the period 2005–2017. CMIP6 ensemble mean maps show projected rate of change 2015–2100 for (d) SSP5-8.5 and (g) SSP1-2.6 scenarios. Also shown are the projected change in 0–700 m OHC for '''(d)''' SSP1-2.6 and '''(g)''' SSP5-8.5 in the CMIP6 ensembles, for the period 2091–2100 versus 2005–2014. 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).

Below 2000 m, direct observations of full-depth ocean temperature change are limited to ship-based, high-quality deep-ocean temperature measurements. Such high-quality full-depth ship-based sampling has improved from 1990 to the present due to the World Ocean Circulation Experiment (WOCE) and the Global Ocean Ship-based Hydrographic Investigations Program (GO-SHIP; [[#Sloyan--2019|Sloyan et al., 2019]] ). The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessed that the ''likely'' warming of the ocean since the 1990s below 2000 m is associated with a marked regional pattern, with larger warming in the Southern Ocean. In the deep North Atlantic, warming has reversed to cooling over the past decade, possibly due to internal variability fed by North Atlantic Deep Water ( [[#9.2.2.3|Section 9.2.2.3]] ). Over the past decade, the warming rate of Antarctic Bottom Water (AABW; [[#9.2.2.3|Section 9.2.2.3]] ) has been dependent on origin: slower from the Weddell Sea and faster from the Ross Sea and Adélie Land. One new study ( [[#Purkey--2019|Purkey et al., 2019]] ) strengthens confidence in AABW warming: below 4000 m a monotonic, basin‐wide, and multi-decadal temperature change is found in the southern Pacific basin, with larger warming rates near the bottom water formation sites than further downstream. New analysis of one model provides ''limited evidence'' that the sparse observational record may underestimate the rate of deep-ocean warming from 1990 to 2010 by about 20% ( [[#Garry--2019|Garry et al., 2019]] ) which is included in the assessed OHC error (Cross-Chapter Box 9.1). There is still ''low agreement'' in deep-ocean changes from ocean data assimilation reanalyses ( [[#Palmer--2017|Palmer et al., 2017]] ) and ''low confidence'' in such inferences. In summary, while observational coverage below 2000 m is sparser than in the upper 2000 m, there is ''high confidence'' that deep-ocean warming below 2000 m has been larger in the Southern Ocean than in other ocean basins due to widespread AABW warming.

Different processes drive OHC patterns over a range of time scales. Recent literature has highlighted the role of ocean circulation variability in driving OHC patterns by decomposing the global pattern of OHC change into a combination of added heat due to climate change taken up under fixed ocean circulation (‘added heat’), and redistribution of heat associated with changing ocean currents (‘redistributed heat’; [[#Gregory--2016|Gregory et al., 2016]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). Redistributed heat alters regional patterns of heat storage and carbon storage (Cross-Chapter Box 5.3; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ; [[#Todd--2020|Todd et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ) but does not affect the global OHC. There is ''medium confidence'' that decadal variability of the ocean circulation strengthened the rate of ocean warming in the Southern Hemisphere compared to the Northern Hemisphere in the decade from 2005 ( [[#Rathore--2020|Rathore et al., 2020]] ; L. [[#Wang--2021|]] [[#Wang--2021|Wang et al., 2021]] ; [[#Zika--2021|Zika et al., 2021]] ). More generally, since 2005, the OHC pattern observed is predominantly due to heat redistribution with regions of both warming and cooling (Figure 9.6; [[#Zika--2021|Zika et al., 2021]] ); however, extending analysis back to 1972 shows the importance of added heat setting a large-scale warming pattern with mid-latitude maxima consistent with subduction of water masses, particularly in Southern Hemisphere Mode Waters ( [[#9.2.2.3|Section 9.2.2.3]] , and Figures 9.6 and 9.8; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). The longer the analysis window, the more added heat dominates over redistributed heat. This translates into more ocean area with statistically significant warming trends and less area with statistically significant cooling trends ( [[#Johnson--2020|Johnson and Lyman, 2020]] ). The region where added heat is most compensated for by redistributed cooling is in the northern North Atlantic basin, where changes in the subpolar gyre circulation and Atlantic Meridional Overturning Circulation (AMOC) result in cooling ( [[#9.2.3.1|Section 9.2.3.1]] ; [[#Williams--2015|]] [[#Williams--2015|Williams et al., 2015]] ; [[#Piecuch--2017|Piecuch et al., 2017]] ; [[#Zanna--2019|Zanna et al., 2019]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). In summary, and strengthening SROCC assessment, ocean warming is not globally uniform due to patterns of uptake predominantly along known water mass pathways, and due to changing ocean circulation redistributing heat within the ocean ( ''high confidence'' ).

While heat redistribution reflects changes in ocean circulation and is a useful concept to understand the underlying processes driving OHC patterns, change in ocean heat transport (OHT) arises due to changes in ocean circulation and ocean temperature and affects regional OHC change. The AR5 did not assess change in OHT and SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) only assessed projected OHT increases into the Nordic Seas and the Arctic Ocean. New evidence of increasing northward OHT into the Arctic has been observed in recent decades ( [[#Muilwijk--2018|Muilwijk et al., 2018]] ; [[#Wang--2019|Q. Wang et al., 2019]] ; [[#Tsubouchi--2021|Tsubouchi et al., 2021]] ), similar to SROCC assessment, and consistent with observed increase in OHC in the ice-free Arctic ocean ( [[#Mayer--2019|Mayer et al., 2019]] ). It is estimated that an increase of 0.021 PW of OHT occurred after 2001 into the Arctic, which is sufficient to account for the recent OHC change in the northern seas ( [[#Tsubouchi--2021|Tsubouchi et al., 2021]] ). However, these trends cannot yet be attributed to anthropogenic forcing due to potential internal variability ( [[#Muilwijk--2018|Muilwijk et al., 2018]] ; [[#Wang--2019|]] [[#Wang--2019|]] [[#Wang--2019|Wang et al., 2019]] ). New evidence strengthens the case that El Niño–Southern Oscillation (ENSO) and the Northern Annular Mode affect interannual OHT variability ( [[#Trenberth--2019|Trenberth et al., 2019]] ) and shows that a slowing AMOC reduces northward OHT in the Atlantic at 26.5°N ( [[#9.2.3.1|Section 9.2.3.1]] and Figure 9.8; [[#Bryden--2020|Bryden et al., 2020]] ). Despite a decrease of AMOC northward heat (0.17 PW) and mass (2.5 Sverdrup (Sv); 1 Sv = 10 <sup>9</sup> kg s <sup>–1</sup> ) transport, OHT has increased toward the Arctic through increased upper northern North Atlantic temperatures and stronger wind-driven gyres ( ''medium confidence'' ) ( [[#9.2.3.4|Section 9.2.3.4]] and Figure 9.11; [[#Singh--2017|Singh et al., 2017]] ; [[#Oldenburg--2018|Oldenburg et al., 2018]] ). In summary, OHT has increased toward the Arctic in recent decades, which at least partially explains the recent OHC change in the Arctic ( ''medium confidence'' ).

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[[File:ee36579304184737ff8bf4ac205e04df IPCC_AR6_WGI_Figure_9_7.png]]

'''Figure''' '''9.7 |''' '''Meridional-depth profiles of zonal-mean potential temperature in the ocean and its rate of change in the upper 2000 m of the Global, Pacific, Atlantic and Indian oceans.''' Shown are '''(a, e, i, m)''' observed temperature (Argo climatology 2005–2014), '''(b, f, j, n)''' bias of the Coupled Model Intercomparison Project Phase 6 (CMIP6) ensemble over this period, and future changes under '''(c, g, k, o)''' SSP1-2.6 and '''(d, h, l, p)''' SSP5-8.5. 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).

Major volcanic eruptions have caused interannual to decadal cooling phases within the marked long-term increase in global OHC – Mount Agung in 1963, El Chichón in 1982 and Mount Pinatubo in 1991 (Cross-Chapter Box 4.1; [[#Church--2005|Church et al., 2005]] ; [[#Fasullo--2016|Fasullo et al., 2016]] ; [[#Stevenson--2016|Stevenson et al., 2016]] ; [[#Fasullo--2018|Fasullo and Nerem, 2018]] ). In the first few years following an eruption, heat exchange with the subsurface ocean allows atmospheric cooling to be sequestered into the seasonal thermocline, therefore reducing the magnitude of the peak atmospheric temperature anomaly ( [[#Gupta--2018|Gupta and Marshall, 2018]] ). However, while explosive volcanic eruptions only disturb the Earth’s radiative budget and surface fluxes for a few years, the ocean preserves an anomaly in OHC in the upper 500 m (also affecting thermosteric sea level) many years after the eruption ( [[#Gupta--2018|Gupta and Marshall, 2018]] ; [[#Bilbao--2019|Bilbao et al., 2019]] ). The anomaly affects the atmosphere through air–sea heat fluxes with surface conditions returning to normal only after several decades ( [[#Gupta--2018|Gupta and Marshall, 2018]] ; [[#Bilbao--2019|Bilbao et al., 2019]] ), or on centennial time scales in the case of repeated eruptions (G.H. [[#Miller--2012|]] [[#Miller--2012|Miller et al., 2012]] ; [[#Atwood--2016|Atwood et al., 2016]] ; [[#Gupta--2018|Gupta and Marshall, 2018]] ). In summary, there is ''medium confidence'' that oceanic mechanisms buffer the atmospheric response to volcanic eruptions on annual time scales by storing volcanic cooling in the subsurface ocean, affecting OHC and thermosteric sea level on decadal to centennial time scales.

CMIP5 and CMIP6 models simulate OHC changes that are consistent with the updated observational and improved estimates of OHC over the period 1960 to 2018 (Figures 9.6, 9.7 and 9.8), and they replicate the vertical partitioning of OHC change for the industrial era, although with a tendency to underestimate OHC gain shallower than 2000 m and overestimate it deeper than 2000 m ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.1.3|Section 3.5.1.3]] ). The AR5 ( [[#Flato--2013|Flato et al., 2013]] ) assessed that climate models transport heat downward more than the real ocean. Since AR5, studies have shown that increasing the horizontal resolution of ocean models tends to increase agreement of vertical heat transport with observations as the dependency on ad-hoc choices of eddy parametrizations is relaxed ( [[#Griffies--2015|Griffies et al., 2015]] ; [[#Chassignet--2020|Chassignet et al., 2020]] ). The magnitude of the AMOC and Indonesian Throughflow affect future OHC change – for example, through overestimated modelled downward heat pumping ( [[#Kostov--2014|Kostov et al., 2014]] ) – and there are indications of greater model consistency in these transports at higher resolution (Figure 9.10; [[#Chassignet--2020|Chassignet et al., 2020]] ; [[#Jackson--2020|]] [[#Jackson--2020|L.C. Jackson et al., 2020]] ). Climate models tend to reproduce the observed added heat, but redistributed heat is less well represented (Figure 9.8; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ; [[#Dias--2020|Dias et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). Since redistributed heat dominates historical OHC change, historical simulations poorly reproduce regional patterns, but as future OHC change will become dominated by added heat, more skill in future modelled OHC patterns is expected ( [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). In summary, climate models have more skill in representing OHC change from added heat than from ocean circulation change ( ''high confidence'' ). Since added heat dominates over redistributed heat on a centennial scale (especially under high-emissions scenarios) confidence in future modelled OHC patterns at the end of the 21st century is greater than at decadal scale.

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[[File:33575866d44b94baa6c55276cf5ddc36 IPCC_AR6_WGI_Figure_9_8.png]]

'''Figure''' '''9.8 |''' '''Decomposition of simulated ocean heat content and northward ocean heat transport. (a, c, e)''' Total ocean heat content (0–2000 m) warming rate as observed and simulated by Coupled Model Intercomparison Project Phase 5 (CMIP5) models over the historical period (1972–2011) and under the RCP8.5 future (2021–2060) versus the associated decomposed '''(b, d, f)''' added heat contribution (neglecting changes in ocean circulation) to the total ( [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ). '''(g)''' Relationship between northward heat transport and Atlantic Meridional Overturning Circulation (AMOC) in HighResMIP models (1950–2050) and observations during the RAPID period (2004–2018). Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessed that the ocean will continue to take up heat in the coming decades for all plausible scenarios, and here this assessment is confirmed with ''very high confidence'' . The SROCC reported that, compared with the observed changes since the 1970s, the warming of the ocean by 2100 would ''very likely'' double to quadruple for low-emissions scenarios (RCP2.6) and increase five to seven times for high-emissions scenarios (RCP8.5). The SROCC also concluded with ''high confidence'' that the overall warming of the ocean would continue this century, even after radiative forcing and mean surface temperatures stabilize. The SROCC projected that OHC in the 0–2000 m layer will increase from 2017 to 2100 by 0.900 ± 0.345 YJ (1 YJ = 10 <sup>24</sup> Joules) under RCP2.6 and 2.150 ± 0.540 YJ under RCP8.5. Updating SROCC estimates with CMIP6 projections gives heat content increases and 17–83% ranges in the 0–2000 m layer between 1995–2014 and 2081–2100 of 1.06 (0.80–1.31) YJ, 1.35 (1.08–1.67) YJ, 1.62 (1.37–1.91) YJ, 1.89 (1.60–2.29) YJ under scenarios SSP1-2.6, SSP2-4.5, SSP3-7.0, and SSP5-8.5, respectively (Figure 9.6 and Table 9.1). The two-layer model used here to calculate thermosteric sea level rise (9.SM.4) and tuned for AR6-assessed equilibrium climate sensitivity (ECS; Section 7.SM.2), provides consistent 17–83% ranges of 1.18 (0.99–1.42) YJ, 1.56 (1.33–1.86) YJ, 1.90 (1.63–2.21) YJ, 2.23 (1.92–2.64) YJ under scenarios SSP1-2.6, SSP2-4.5, SSP3-7.0, and SSP5-8.5, respectively (Table 9.1). Based on CMIP6 models and the two-layer model, it is ''likely'' that, between 1995–2014 and 2081–2100, OHC will increase two to four times the amount of the 1971–2018 OHC increase under SSP1-2.6, and four to eight times that amount under SSP5-8.5. The CMIP6 models show that OHC dependence on scenarios begins only after about 2040 (Figure 9.6).

The OHC patterns projected by CMIP6 models (Figures 9.6 and 9.7) are similar to the CMIP5 projections assessed in SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ): faster warming in all water mass subduction regions (e.g., subtropical cells and mode waters); deeper penetration in the centre of subtropical gyres; slower northern North Atlantic warming due to slowing AMOC; and slower subpolar Southern Ocean warming due upwelled pre-industrial water masses. Decreased aerosol forcing will allow Northern Hemisphere ocean warming to be faster and less dominated by Southern Hemisphere change ( [[#Shi--2018|Shi et al., 2018]] ; [[#Irving--2019|Irving et al., 2019]] ). Since SROCC, distinguishing between added and redistributed heat has aided in understanding projections ( [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ; [[#Dias--2020|Dias et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). The near-term decades will feature patterns strongly influenced by heat redistribution and internal variability ( [[#Rathore--2020|Rathore et al., 2020]] ). Strengthening Southern Hemisphere westerlies are projected, except for stringent mitigation scenarios ( [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ), and will cause a northward and downward OHT. There is ''low agreement'' in future Southern Ocean warming across model results due to uncertainties in the magnitude of westerly wind changes (Figure 9.4; [[#Liu--2018|Liu et al., 2018]] ; [[#He--2019|He et al., 2019]] ; [[#Dias--2020|Dias et al., 2020]] ; [[#Lyu--2020b|Lyu et al., 2020b]] ) and the degree of eddy compensation of overturning across different parametrizations and resolutions ( [[#9.2.3.2|Section 9.2.3.2]] ; [[#Beal--2016|Beal and Elipot, 2016]] ; [[#Mak--2017|Mak et al., 2017]] ; [[#Roberts--2020|Roberts et al., 2020]] ). By 2100, however, the OHC change will be dominated by the added heat response, particularly for strong warming scenarios ( [[#Garuba--2018|Garuba and Klinger, 2018]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ) with added heat following unperturbed water mass pathways in the North Atlantic and Southern Ocean (Figure 9.8; [[#Dias--2020|Dias et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). There is ''high confidence'' that projected weakening of the AMOC ( [[#9.2.3.1|Section 9.2.3.1]] ) will cause a decrease in northward OHT in the Northern Hemisphere mid-latitudes (Figure 9.8 and Sections 9.2.3.1 and 4.3.2.3; [[#Weijer--2020|Weijer et al., 2020]] ) associated with a dipole pattern of Atlantic OHC redistributed from northern to low latitudes that may override added heating in the northern North Atlantic (Figures 9.6, 9.7 and 9.8). Variations in the degree of AMOC redistributed heat ( [[#Menary--2018|Menary and Wood, 2018]] ) causes large intermodel spread in SST (Figure 9.3) and OHC change (Figure 9.6; [[#Kostov--2014|Kostov et al., 2014]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ; [[#Todd--2020|Todd et al., 2020]] ; [[#Couldrey--2021|Couldrey et al., 2021]] ). In the 700–2000 m depth range, CMIP5 and CMIP6 models project the largest warming to be in the North Atlantic Deep Water and Antarctic Intermediate Water (Figure 9.7) while below 2000 m, the North Atlantic cools in many models, and Antarctic Bottom Waters warm ( [[#Sallée--2013b|Sallée et al., 2013b]] ; [[#Heuzé--2015|Heuzé et al., 2015]] ). In summary, on decadal time scales, redistribution will dominate regional patterns of OHC change without affecting the globally integrated OHC; however, by 2100, particularly under strong warming scenarios, there is ''high confidence'' that regional patterns of OHC change will be dominated by added heat entering the sea surface, primarily in water mass formation regions in the subtropics; and reduced aerosols will increase the relative rate of Northern Hemisphere heat uptake ( ''medium confidence'' ).

The SROCC assessed that the warming of the deep ocean is slow to manifest, with multi-century or longer response times, so global OHC (and global mean thermosteric sea level) will continue to rise for centuries (Figures 9.9 and 9.30). New studies show that this continuation persists, even after cessation of greenhouse gas emissions ( [[#Ehlert--2018|Ehlert and Zickfeld, 2018]] ). Ocean warming will continue, even after emissions reach zero because of slow ocean circulation ( [[#Larson--2020|Larson et al., 2020]] ). OHC will increase until at least 2300, even for low-emissions scenarios, but with a scenario-dependent rate ( [[#Nauels--2017|Nauels et al., 2017]] ; [[#Palmer--2018|Palmer et al., 2018]] ) and depends on cumulative CO <sub>2</sub> emissions, as well as the time profile of emissions ( [[#Bouttes--2013|Bouttes et al., 2013]] ). Past long-term changes in total OHC illustrate adjustment relevant to expected future changes (Figure 9.9). Observational data from ice core rare gas elemental and isotopic ratios document a rise in global OHC relative to the Last Glacial Maximum of &gt;17,000 ZJ (change in mean ocean temperature &gt;3.1°C; 1 ZJ = 10 <sup>21</sup> Joules) (Figure 9.9; [[#Bereiter--2018|Bereiter et al., 2018]] ; [[#Baggenstos--2019|Baggenstos et al., 2019]] ; [[#Shackleton--2019|Shackleton et al., 2019]] , 2020). This temperature increase is significantly larger than the modelled OHC changes associated with collapse of AMOC alone, and tracks rising Southern Ocean SST ( [[#Uemura--2018|Uemura et al., 2018]] ), strengthening of the deep abyssal overturning cell ( [[#Du--2020|Du et al., 2020]] ) and increased North Atlantic water in the Southern Ocean ( [[#Wilson--2020|Wilson et al., 2020]] ). This underscores the importance of Antarctic abyssal ventilation on long-term oceanic heat budgets ( [[#9.2.3.2|Section 9.2.3.2]] ). An ensemble of four intermediate-complexity models project 10,000-year future responses to CO <sub>2</sub> emissions ( [[#Clark--2016|Clark et al., 2016]] ) with SST change peaking around 2300 and a varying scenario-dependent magnitude approaching the scale of glacial-to-interglacial changes in paleodata (Figure 9.9). Long-term OHC commitments relative to 1850–1900 conditions are 2.6, 9.7, 15.2, 21.6, and 28.0 YJ (with mean ocean temperature change as much as 5.1°C) for emissions of 0, 1280, 2560, and 3840 and 5120 Gt after 2000 CE respectively, with OHC peaking near 4000 CE, reflecting whole-ocean warming lagging SST by thousands of years. The exact timing is uncertain, subject to rates of high-latitude meltwater input ( [[#Van%20Breedam--2020|Van Breedam et al., 2020]] ) and circulation time ( [[#Gebbie--2019|Gebbie and Huybers, 2019]] ). In summary, there is ''very'' ''high confidence'' that there is a long-term commitment to increased OHC in response to anthropogenic CO <sub>2</sub> emissions, which is essentially irreversible on human time scales.

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[[File:ac6315ed3e1fc01ef20199f8b35b8d18 IPCC_AR6_WGI_Figure_9_9.png]]

'''Figure 9.9 |''' '''Long-term trends of ocean heat content (OHC) and surface temperature. (a, b)''' Ice-core rare gas estimates of past mean OHC (ZJ), scaled to global mean ocean temperature (°C), and to steric global mean sea level (GMSL) (m) per CCB-2 (red dashed line), compared to surface temperatures (black solid line, gold solid line; °C rightmost axis). Southern Ocean sea surface temperature (SST) from multiple proxies in 11 sediment cores and from ice core deuterium excess ( [[#Uemura--2018|Uemura et al., 2018]] ). '''(a)''' Penultimate glacial interval to last interglacial, 150,000–100,000 yr B2K (before 2000) ( [[#Shackleton--2020|Shackleton et al., 2020]] ). '''(b)''' Last glacial interval to modern interglacial, 40,000–0 yr B2K ( [[#Baggenstos--2019|Baggenstos et al., 2019]] ; [[#Shackleton--2019|Shackleton et al., 2019]] ). Changes in OHC (dashed lines) track changes in Southern Ocean SST (solid lines). '''(c)''' Long-term projected (2000 to 12000 CE) changes of OHC (dashed lines) in response to four greenhouse gas emissions scenarios ( [[#Clark--2016|Clark et al., 2016]] ) scale similarly to large-scale paleo changes but lag projected global mean SST (solid lines). '''(d)''' model simulated 1500–1999 OHC ( [[#Gregory--2006|Gregory et al., 2006]] ) and 1955–2019 observations ( [[#Levitus--2012|Levitus et al., 2012]] ) updated by NOAA NODC. All data expressed as anomalies relative to pre-industrial time. Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

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<span id="ocean-salinity"></span>
==== 9.2.2.2 Ocean Salinity ====

<div id="h3-5-siblings" class="h3-siblings"></div>

The AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ) assessed that it was ''very likely'' that subsurface salinity changes reflect surface salinity change, and that basin-scale regions of high salinity and evaporation had trended more saline, while regions of low salinity and more precipitation had trended fresher since the 1950s. The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessment was consistent with AR5. [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.2|Section 2.3.3.2]] strengthens evidence that subsurface salinity trends are connected to surface trends ( ''very likely'' ) , which are, in turn, linked to an intensifying hydrological cycle ( ''medium confidence'' ). Increasing evidence from updated observational records indicates that it is now ''virtually certain'' that surface salinity contrasts are increasing. At basin scale, [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.2|Section 2.3.3.2]] and AR5 concur that it is ''very likely'' that the Pacific and Southern Ocean have freshened, and the Atlantic has become more saline. Figures 3.25 and 3.27 compare CMIP6 models to salinity observations.

Globally the mean salinity contrast at near-surface between high- and low-salinity regions increased 0.14 [0.07 to 0.20] from 1950 to 2019 ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.2|Section 2.3.3.2]] ). At regional scale, SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) assessed an Arctic liquid freshwater trend of 600 ± 300 km <sup>3</sup> yr <sup>–1</sup> (600 ± 200 Gt yr <sup>–1</sup> ) between 1992 and 2012, reflecting changes associated with continental freshwater imports that affect ocean mass (land ice, rivers) as well as changes in sea ice volume. Since AR5, regional observation-based analyses not assessed in SROCC further confirm the long-term, large-scale and regional patterns of salinity change, both at the ocean surface and in the subsurface ocean, including almost 120 years of changes in the North Atlantic ( [[#Friedman--2017|Friedman et al., 2017]] ) and 60 years of monitoring in the subpolar North Pacific ( [[#Cummins--2020|Cummins and Ross, 2020]] ). These longer time series also provide context to detect large multi-annual change from 2012 to 2016 in the subpolar North Atlantic, unprecedented over the centennial record ( [[#Holliday--2020|Holliday et al., 2020]] ). In summary, there is ''high confidence'' that salinity trends have extended for more than 60 to 100 years in the regions with long historical observation records, such as the North Pacific and the North Atlantic basin.

While there is ''low confidence'' in direct estimates of trends in surface freshwater fluxes (Sections 2.3.1.3.5, 8.3.1.1 and 9.2.1.2), as discussed in SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ), observational studies coupled with modelling studies suggest that surface flux changes drive many observed near-surface salinity changes, on top of changes specific to polar regions. Advances in salinity observations – for example, the Argo program ( [[#Riser--2016|Riser et al., 2016]] ); Soil Moisture and Ocean Salinity (SMOS), Aquarius and Soil Moisture Active Passive (SMAP; [[#Supply--2018|Supply et al., 2018]] ; [[#Vinogradova--2019|Vinogradova et al., 2019]] ) – combined with process studies (SPURS-1/2; [[#Lindstrom--2015|Lindstrom et al., 2015]] ; SPURS-2 Planning Group 2015) and methodological and numerical advances, have increased understanding of how subsurface salinity anomalies link to surface fluxes, and thus increase confidence that near-surface and subsurface salinity pattern changes since the 1950s are linked to changing surface freshwater fluxes ( [[#Zika--2018|Zika et al., 2018]] ; [[#Cheng--2020|Cheng et al., 2020]] ) with an additional contribution from changes in sea ice and land ice discharge at high latitudes ( [[#Haumann--2016|Haumann et al., 2016]] ; [[#Purich--2018|Purich et al., 2018]] ; [[#Dukhovskoy--2019|Dukhovskoy et al., 2019]] ; [[#Rye--2020|Rye et al., 2020]] ). There is therefore ''medium confidence'' in the processes linking surface fluxes to surface and subsurface salinity change.

Ocean circulation changes also affect salinity, largely on annual to decadal time scales ( [[#Du--2019|Du et al., 2019]] ; [[#Liu--2019|Liu et al., 2019]] ; [[#Holliday--2020|Holliday et al., 2020]] ). For instance, in the subpolar North Atlantic, increasing northward transport of Atlantic waters entering the subpolar gyre from the South have compensated the salinity decrease expected from increased Greenland meltwater flux since the early 1990s ( [[#Dukhovskoy--2016|Dukhovskoy et al., 2016]] , 2019; [[#Stendardo--2020|Stendardo et al., 2020]] ). After the mid-2010s the trend reversed towards a broad freshening, the largest in 120 years, in the North Atlantic ( [[#Holliday--2020|Holliday et al., 2020]] ). The long-term freshening in the Pacific Ocean has also been subject to decadal variability, such as a marked salinification since 2005 associated with increased surface fluxes (G. [[#Li--2019|]] [[#Li--2019|]] [[#Li--2019|Li et al., 2019]] ). Local salinity anomalies forced by water cycle intensification can be weakened by rapid exchange between basins with opposing trends, such as by water mass exchange in shallow wind-driven cells between the tropics and the subtropics ( [[#Levang--2020|Levang and Schmitt, 2020]] ). Similarly, eddy exchanges between neighbouring gyres can partly counterbalance decadal time scale long-term subpolar freshening and affect deep convection ( [[#Levang--2020|Levang and Schmitt, 2020]] ). There is ''high confidence'' that, at annual to decadal time scales, regional salinity changes are driven by ocean circulation change superimposed on longer-term trends.

The CMIP5 historical simulations have patterns similar to, but with greater spatial variability than, observed estimates and correspondingly smaller amplitudes in the multi-model mean ( [[#Durack--2015|Durack, 2015]] ; [[#Cheng--2020|Cheng et al., 2020]] ; [[#Silvy--2020|Silvy et al., 2020]] ). [[IPCC:Wg1:Chapter:Chapter-3#3.5.2.1|Section 3.5.2.1]] reports, however, that the fidelity of ocean salinity simulation has improved in CMIP6, and near-surface and subsurface biases have been reduced ( ''medium confidence'' ), though the structure of the biases strongly reflects those of CMIP5. At regional scale, salinity biases are at least partially a result of inaccurate ocean dynamics ( [[#Levang--2020|Levang and Schmitt, 2020]] ). Despite the regional limitations, [[IPCC:Wg1:Chapter:Chapter-3#3.5.2.2|Section 3.5.2.2]] assesses that, at the global scale, it is ''extremely likely'' that human influence has contributed to observed surface and subsurface salinity changes since the mid-20th century (strengthened from the ''very likely'' AR5 assessment).

The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessed that projected salinity changes in the subsurface ocean reflect changes in the rates of formation of water masses or their newly formed properties. Additional consistent newer evidence based on CMIP5 and regional climate models confirms that 21st century projections adhere to the ‘fresh gets fresher, salty gets saltier’ paradigm, through subduction of freshening high-latitude waters into the ventilated water masses in both hemispheres in the Pacific, Indian and Southern Ocean – especially the Arctic and upper Southern Ocean, and saltier subtropical and Mediterranean surface waters – lead to saltier pycnoclines and North Atlantic mode water ( [[#Metzner--2020|Metzner et al., 2020]] ; [[#Parras-Berrocal--2020|Parras-Berrocal et al., 2020]] ; [[#Silvy--2020|Silvy et al., 2020]] ; [[#Soto-Navarro--2020|Soto-Navarro et al., 2020]] ). Overall, projections confirm SROCC assessment that fresh ocean regions will continue to get fresher and salty ocean regions will continue to get saltier in the 21st century ( ''medium confidence'' ).

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<span id="water-masses"></span>
==== 9.2.2.3 Water Masses ====

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Water masses refer to connected bodies of ocean water, formed at the ocean surface with identifiable properties (temperature, salinity, density, chemical tracers) resulting from the unique formation conditions of the overlying atmosphere and/or ice, before being transferred (subducted) to the deeper ocean below the surface turbulent layer. As water masses subduct, they ventilate the subsurface ocean, transferring characteristics acquired at the ocean surface to the subsurface. By integrating surface flux changes, water masses provide higher signal-to-noise ratios for detecting and monitoring climate change than surface fluxes ( [[#Bindoff--2000|Bindoff and McDougall, 2000]] ; [[#Durack--2010|Durack and Wijffels, 2010]] ; [[#Silvy--2020|Silvy et al., 2020]] ).

Subtropical mode waters (STMW) ventilate the main thermocline of the ocean at mid- to low-latitudes and have circulation time scales away from the surface of the order of years to decades. The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) reported that warming in the subtropical gyres penetrates deeper than in other gyres, following the density surfaces in these gyres. Consistently, we assess that STMW have deepened worldwide, with greatest deepening in the Southern Hemisphere ( ''high confidence'' ) ( [[#Häkkinen--2016|Häkkinen et al., 2016]] ; [[#Desbruyères--2017|Desbruyères et al., 2017]] ). Subsurface warming in the Northern Hemisphere STMW is larger than at the surface ( [[#Sugimoto--2017|Sugimoto et al., 2017]] ) because they are formed in winter western boundary current extensions, where surface warming is larger than the global average ( [[#9.2.1.1|Section 9.2.1.1]] ). Variability in STMW thickness or temperature has a large imprint on OHC ( [[#9.2.2.1|Section 9.2.2.1]] ; [[#Kolodziejczyk--2019|Kolodziejczyk et al., 2019]] ). STMW are observed to be freshening in the North Pacific and associated with increased salinity in the North Atlantic ( [[#Oka--2017|Oka et al., 2017]] ; [[#Silvy--2020|Silvy et al., 2020]] ), with large decadal variability ( [[#Oka--2019|Oka et al., 2019]] ; [[#Wu--2020|Wu et al., 2020]] ). Anthropogenic temperature and salinity changes in the STMW layer are projected to intensify in the future, with emergence from natural variability around 2020 to 2040 ( [[#Silvy--2020|Silvy et al., 2020]] ).

Subantarctic mode water (SAMW) and Antarctic intermediate water (AAIW) form at the Southern Ocean surface directly north of the Antarctic Circumpolar Current and ventilate the upper 1000 m of the Southern Hemisphere subtropics. The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) reported a freshening of these water masses between 1950 and 2018, and they are projected to have the largest subsurface temperature increase of the Southern Hemisphere oceans, along with a continued freshening, in the 21st century. The SROCC connected SAMW and AAIW to Southern Ocean temperature changes as the large Southern Ocean surface heat uptake is circulated and mixed along with these water masses ( ''high confidence'' ). Close to its formation region, SAMW is predominantly affected by air–sea flux changes, while further northward it is influenced by wind-forced changes ( [[#Meredith--2019|Meredith et al., 2019]] ). New evidence shows that a change in SAMW heat content over the last decade is primarily attributable to its thickening ( [[#Kolodziejczyk--2019|Kolodziejczyk et al., 2019]] ). Over the past decade, the SAMW and AAIW volumes have changed by thickening of the lighter and thinning of the denser parts of SAMW and AAIW, leading to lightening of these ventilated ocean layers overall ( [[#Hong--2020|Hong et al., 2020]] ; [[#Portela--2020|Portela et al., 2020]] ). Over the last decade, there is ''limited evidence'' of increased subduction of SAMW due to deepening mixed layers in the SAMW formation region ( [[#9.2.1.3|Section 9.2.1.3]] ; [[#Qu--2020|Qu et al., 2020]] ). Climate models from CMIP3 to CMIP5 generally simulated shallower and lighter SAMW and AAIW than is observed ( [[#Flato--2013|Flato et al., 2013]] ). New analysis of CMIP5 models suggests that the freshening of these water masses is one of the most prominent projected salinity changes in the world ocean, and that this freshening emerged from internal variability as early as the 1980s to 1990s ( [[#Silvy--2020|Silvy et al., 2020]] ).

Trends in North Atlantic Deep Water (NADW) are obscured by decadal variability ( [[#Rhein--2013|Rhein et al., 2013]] ; [[#Bindoff--2019|Bindoff et al., 2019]] ). The AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ) assessed that it is ''very likely'' that the temperature, salinity, and formation rate of the Upper NADW (formed by deep convection in the Labrador and Irminger Seas) is dominated by strong decadal variability related to the North Atlantic Oscillation (NAO) and it is ''likely'' that Lower NADW (formed in the Nordic Seas and supplied to the North Atlantic by deep overflows over the sills between Scotland and Greenland) cooled from 1955 to 2005. New insights from observations have emphasized the stability of the deep overflows associated with Lower NADW ( [[#Hansen--2016|Hansen et al., 2016]] ; [[#Jochumsen--2017|Jochumsen et al., 2017]] ; [[#Østerhus--2019|Østerhus et al., 2019]] ) and even slight warming in the Faroe Bank Channel ( [[#Hansen--2016|Hansen et al., 2016]] ). As a result, the AR5 assessment that Lower NADW ''likely'' cooled between 1955 and 2005 is revised to: it is ''likely'' that any observed changes in temperature, salinity, and formation rate of the Lower NADW are dominated by decadal variability. For CMIP5 models, it was shown that AMOC variability is linked to variability in NADW formation ( [[#Heuzé--2017|Heuzé, 2017]] ) and projected AMOC decline to decreased NADW formation (both Lower NADW and Upper NADW; [[#Heuzé--2015|Heuzé et al., 2015]] ). For CMIP6 models, projected AMOC decline is also associated with a decline in NADW formation ( [[#Reintges--2017|Reintges et al., 2017]] ; [[#Weijer--2020|Weijer et al., 2020]] ). The link between AMOC and NADW formation appears insensitive to the large range in model bias in NADW water mass characteristics ( [[#Heuzé--2017|Heuzé, 2017]] ). Many models may overestimate deep water formation in the Labrador Sea, but at least one new model is consistent with recent Overturning in the Subpolar North Atlantic Program (OSNAP) observations showing very weak overturning in the western subpolar gyre, where Labrador Sea water is formed ( [[#Menary--2020a|Menary et al., 2020a]] ). The CMIP6 models show a reduced bias in NADW properties compared to CMIP5 models, but still feature varying locations of deep convection in the subpolar gyre: some convect only in the Labrador Sea (6/35 models), most in both the Labrador and Irminger Seas (26/35 models; as is observed), and some only in the Irminger Sea (3/35 models), but in general, the area where deep convection takes place has expanded relative to CMIP5, which appears unrealistic ( [[#Heuzé--2021|Heuzé, 2021]] ). Models with most deep convection in the subpolar gyre feature the smallest bias in NADW characteristics, partly associated with NADW formed in the Nordic Seas (as observed) being largely unable to leave the area ( [[#Heuzé--2021|Heuzé, 2021]] ) due to inaccurate overflows ( [[#Danabasoglu--2010|Danabasoglu et al., 2010]] ; [[#Deshayes--2014|Deshayes et al., 2014]] ; [[#Wang--2015|]] [[#Wang--2015|Wang et al., 2015]] ). Despite the wide range in model bias, it remains ''very likely'' that any long-term (multi-decadal or longer) decrease in AMOC is accompanied by a decline in NADW formation, associated with lighter densities in the northern North Atlantic and Arctic basins.

The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) assessed that the global volume of Antarctic Bottom Water (AABW) had decreased and warmed since the 1980s, most noticeably near Antarctica. The SROCC also noted freshening in the Indian and Pacific sectors of the Southern Ocean and a higher rate of freshening in the Indian Sector from the 2000s to 2010s than from the 1990s to 2000s ( ''low confidence'' ). Since SROCC, freshening of Indian Ocean AABW from 1974 to 2016 has been revealed ( [[#Aoki--2020|Aoki et al., 2020]] ). Additionally, interannual to decadal variability in AABW has been quantified to be larger than previously thought in terms of temperature, salinity and thickness, and in volume transport ( [[#Abrahamsen--2019|Abrahamsen et al., 2019]] ; [[#Purkey--2019|Purkey et al., 2019]] ; [[#Gordon--2020|Gordon et al., 2020]] ; [[#Silvano--2020|Silvano et al., 2020]] ). Multi-decadal to centennial modes of variability could have driven the observed trends of the lower cell over the past decades via the opening of a Weddell Sea Polynya (L. [[#Zhang--2019|]] [[#Zhang--2019|]] [[#Zhang--2019|Zhang et al., 2019]] ), although other studies find it contributed minimally to the observed abyssal warming ( [[#Zanowski--2015|Zanowski et al., 2015]] ; [[#Zanowski--2017|Zanowski and Hallberg, 2017]] ). Therefore, there is ''limited evidence'' and ''low agreement'' in the role of open ocean polynyas in driving past decadal observed trends of AABW. Beyond variability, all observational, theoretical, and numerical evidence supports SROCC assessment that formation and export of AABW will continue to decrease due to warming and freshening of surface source waters near the Antarctic continent. Consistent with [[#9.2.3.2|Section 9.2.3.2]] , confidence in this assessment is increased to ''medium confidence'' compared to SROCC.

Circumpolar Deep Water (CDW) lies in the Southern Ocean and forms by the mixing of NADW and AABW ( [[#Talley--2013|Talley, 2013]] ). The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) assessed with ''low confidence'' that mean southward and upward CDW transport is linked to decadal wind variability ( [[#9.2.3.2|Section 9.2.3.2]] ), and that CDW has warmed south of the Antarctic Circumpolar Current (ACC) in the past decades. New evidence reinforces SROCC assessment: changes in Southern Ocean wind stress have been confirmed to drive variability and increase the large-scale southward CDW transport ( [[#Waugh--2019|Waugh et al., 2019]] ). In addition, growing evidence suggests that the upper-ocean stratification increase in the subpolar Southern Ocean since the 1970s ( [[#9.2.1.3|Section 9.2.1.3]] ) has reduced the volume of CDW that is mixed to the surface, causing subsurface CDW warming ( [[#Bronselaer--2020|Bronselaer et al., 2020]] ; [[#Haumann--2020|Haumann et al., 2020]] ; [[#Jeong--2020|Jeong et al., 2020]] ; [[#Moorman--2020|Moorman et al., 2020]] ). Large regions of the Antarctic shelves are currently isolated from warm CDW ( [[#Thompson--2018|Thompson et al., 2018]] ; [[#Jourdain--2020|Jourdain et al., 2020]] ). The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) assessed that subsurface warming extends close to Antarctica and has co-occurred with shoaling of the CDW since the 1980s, influencing the continental shelf most in the Amundsen-Bellingshausen Seas, Wilkes Land, and the Antarctic Peninsula. New evidence since SROCC reinforces confidence in the importance of the role of winds in transporting heat associated with CDW to continental shelves and ice cavities in the Amundsen-Bellingshausen Seas ( [[#Dotto--2019|Dotto et al., 2019]] ) and via variable small-scale undercurrents to the Shirase Glacier Tongue in East Antarctica ( [[#Hirano--2020|Hirano et al., 2020]] ; [[#Kusahara--2021|Kusahara et al., 2021]] ). There is ''limited evidence'' that increased greenhouse gas forcing has caused a slight mean change of the local winds from 1920 to 2018, facilitating CDW heat intrusion onto the Amundsen-Bellingshausen continental shelf and ice shelf melt ( [[#Holland--2019|Holland et al., 2019]] ). Multiple lines of observational, numerical, theoretical, and paleo evidence provide ''high confidence'' that changes in wind pattern ( [[#Spence--2014|Spence et al., 2014]] ; [[#Dotto--2019|Dotto et al., 2019]] ; [[#Holland--2019|Holland et al., 2019]] ), increased ice-shelf melt ( [[#Golledge--2019|Golledge et al., 2019]] ; [[#Moorman--2020|Moorman et al., 2020]] ), reduction in sea ice production ( [[#Timmermann--2013|Timmermann and Hellmer, 2013]] ; [[#Obase--2017|Obase et al., 2017]] ), and eddies ( [[#Stewart--2015|Stewart and Thompson, 2015]] ; [[#Thompson--2018|Thompson et al., 2018]] ) can facilitate access of CDW to the sub-ice-shelf cavities ( [[#9.4.2.1|Section 9.4.2.1]] ). However, there is ''low confidence'' in the quantitification, importance and the ability of present models, especially at coarse resolution, to project changes in each of these processes ( [[#9.4.2.2|Section 9.4.2.2]] ). Some studies have projected a possible shift from cold to warm sub-ice-shelf cavities causing a sudden flush of warm water underneath ice shelves, but there is ''low confidence'' in the driving processes and the threshold to trigger the shift (Box 9.4; [[#Hellmer--2012|Hellmer et al., 2012]] , 2017; [[#Silvano--2018|Silvano et al., 2018]] ; [[#Hazel--2020|Hazel and Stewart, 2020]] ).

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=== 9.2.3 Regional Ocean Circulation ===

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==== 9.2.3.1 Atlantic Meridional Overturning Circulation ====

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Atlantic Meridional Overturning Circulation (AMOC) is the main overturning current system in the South and North Atlantic oceans. It transports warm upper-ocean water northwards, and cold, deep water southwards, as part of the global ocean circulation system ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4.1|Section 2.3.3.4.1]] ). Changes in AMOC influence global ocean heat content (OHC) and transport ( [[#9.2.2.1|Section 9.2.2.1]] ); global ocean anthropogenic carbon uptake changes and climate sensitivity (Cross-Chapter Box 5.3); and dynamical sea level change ( [[#9.2.4|Section 9.2.4]] ). Since AR5/SROCC, confidence in modelled and reconstructed AMOC has decreased due to new observations and model disagreement. Confidence levels have been revisited in modelled AMOC evolution during the 20th century, the magnitude of 21st-century AMOC decline, and the possibility of an abrupt collapse before 2100.

The AR5 ( [[#Flato--2013|Flato et al., 2013]] ) found that the mean AMOC strength in CMIP5 models ranges from 15 to 30 Sv for the historical period. The multi-model mean overturning at 26°N in CMIP5 and CMIP6 is comparable to the RAPID array measurements ( [[#Reintges--2017|Reintges et al., 2017]] ), but the inter-model spread in CMIP6 is as large (10–31 Sv) as in CMIP5 ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.4|Section 3.5.4]] ; [[#Weijer--2020|Weijer et al., 2020]] ). Biases in simulations of the present-day AMOC and associated deep convection in the subpolar gyre and Nordic Seas were large in CMIP5 models, with many models exhibiting ocean convection that is too deep, over too large an area, too far south, and occurring too frequently ( [[#9.2.1.3|Section 9.2.1.3]] and Figure 9.5; [[#Heuzé--2017|Heuzé, 2017]] ) related to biases in sea ice extent, overflows, and freshwater forcing ( [[#Deshayes--2014|Deshayes et al., 2014]] ; H. [[#Wang--2015|]] [[#Wang--2015|Wang et al., 2015]] ). As a result, the AMOC in CMIP5 was nearly always too shallow, with too weak a temperature contrast between the northward and southward flowing branches. Deep convection errors are still large in CMIP6, and the shallow bias in AMOC persists ( [[#Weijer--2020|Weijer et al., 2020]] ; [[#Heuzé--2021|Heuzé, 2021]] ). Since AR5, there is emerging evidence that enhancing horizontal resolution can reduce long-standing climate model biases in AMOC strength, where the magnitude and profile of northward heat transport at 26°N become more comparable to observations ( [[#Chassignet--2020|Chassignet et al., 2020]] ; [[#Roberts--2020|Roberts et al., 2020]] ). The sensitivity of the AMOC to ocean resolution, however, is model-dependent and can be positive as well as negative ( [[#Roberts--2020|Roberts et al., 2020]] ). An increase in AMOC strength at 26°N, with higher resolution in the ocean component, has been associated with too strong (deep) convection in the subpolar gyre and too deep winter mixed layers ( [[#Jackson--2020|]] [[#Jackson--2020|L.C. Jackson et al., 2020]] ), which occurs in most CMIP6 models that are unable to overflow deep water formed in the Nordic Seas across the Greenland–Iceland–Scotland Ridge. Models with a correct AMOC strength may do so by compensating a lack of deep-water outflow from the Nordic Seas through too much deep convection and deep-water formation in the Labrador and Irminger Seas ( [[#Heuzé--2021|Heuzé, 2021]] ).

Models and paleoreconstructions have often assumed a close relation between the AMOC and deep convection in the Labrador Sea; the Labrador Sea convection variability has been interpreted as connecting to AMOC variability. Observational studies have been inconclusive on whether this relation exists ( [[#Buckley--2016|Buckley and Marshall, 2016]] ). New insight from observed overturning in the eastern and western subpolar gyre in the North Atlantic in OSNAP ( [[#Lozier--2019|Lozier et al., 2019]] ; [[#Petit--2020|Petit et al., 2020]] ) reveals that 15.6 ± 3.1Sv takes place north of the OSNAP array between Greenland and Scotland, with only 2.1 ± 0.9 Sv of overturning occurring across the Labrador Sea, as found with the OSNAP 53°N array spanning the mouth, calling into question the validity of the Labrador Sea convection–AMOC link ( [[#Lozier--2019|Lozier et al., 2019]] ). Although these results are derived from only the first 21 months of data from monitoring since 2014, hydrographic observations during 1990–1997 previously found small overturning (1–2 Sv) in the Labrador Sea ( [[#Pickart--2007|Pickart and Spall, 2007]] ). However, previous estimates of Labrador Sea Water formation (obtained with different techniques) suggest larger overturning ( [[#Haine--2008|Haine et al., 2008]] ). Part of this controversy could be explained if a large fraction of newly formed Labrador Sea Water is not exported from the Labrador Sea. The OSNAP observations are supported by previous hydrographic measurements in showing strong east–west symmetry in isopycnal slope in the Labrador Sea in periods of both strong and weak convection; this implies compensating northward and southward transport above and below the potential density surface that separates the upper and lower overturning limbs ( [[#Lozier--2019|Lozier et al., 2019]] ), despite large deep convection variability ( [[#Yashayaev--2007|Yashayaev, 2007]] ; [[#Yashayaev--2016|Yashayaev and Loder, 2016]] ). New observations of deep winter mixing in the Irminger Basin ( [[#de%20Jong--2018|de Jong et al., 2018]] ; [[#Josey--2019|Josey et al., 2019]] ) support the assertion that the Irminger Sea, in addition to the Nordic Seas ( [[#Chafik--2019|Chafik and Rossby, 2019]] ), are the main sources of overturning in the eastern subpolar gyre, consistent with OSNAP ( [[#Petit--2020|Petit et al., 2020]] ). It is unclear to what extent models are in disagreement with this view of overturning in the subpolar gyre, as a direct comparison with OSNAP of model analyses partitioning the overturning into a western and eastern part is mostly lacking, with a notable exception ( [[#Menary--2020a|Menary et al., 2020a]] ). Other results give rise to considerable uncertainty over veracity of the models in simulating the overturning partitioning between east and west and the role of various drivers of AMOC variability, including: the analysis of water mass formation in CMIP6 models ( [[#Heuzé--2021|Heuzé, 2021]] ); the analysis between Labrador Sea Water formation and AMOC in a suite of ocean-only models ( [[#Danabasoglu--2014|Danabasoglu et al., 2014]] ); and the fact that when the OSNAP observing system design was tested in an eddy-permitting ocean model comparable amounts of overturning in the western and eastern subpolar gyre were found ( [[#Susan%20Lozier--2017|Susan Lozier et al., 2017]] ). Disagreement between models and OSNAP observations may decrease in higher-resolution models ( [[#Menary--2020a|Menary et al., 2020a]] ). In summary, multiple lines of evidence provide ''medium agreement'' between models and observations on drivers of change and variability in the AMOC and, in particular, the role of Labrador Sea deep convection in constituting AMOC variability.

The AMOC is a potential driver of Atlantic Multi-decadal Variability (AMV), but there is new evidence that anthropogenic aerosol changes have contributed to observed AMV changes, and that underestimation of the magnitude and duration of AMV changes in CMIP5 is tempered in CMIP6 ( [[IPCC:Wg1:Chapter:Chapter-3#3.7.7|Section 3.7.7]] and Annex IV.2.7). Comparison of observed AMOC variability at the RAPID section with modelled variability reveals that CMIP5 models appear to largely underestimate the interannual and decadal time scale variability ( [[#Roberts--2014|Roberts et al., 2014]] ; [[#Yan--2018|Yan et al., 2018]] ), and similar results are found when analysing CMIP6 models ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.4.1|Section 3.5.4.1]] ). By underestimating the multi-decadal AMOC–AMV link and other low-frequency AMOC variability, climate models also underestimate internal variability in subpolar SSTs that feed back on the North Atlantic Oscillation (NAO). This causes the NAO to lack variability on multi-decadal time scales ( [[#Kim--2018|Kim et al., 2018]] ). Despite the role of the AMOC in generating AMV through subsurface temperatures in antiphase with SST and downward heat fluxes into the ocean that anticorrelate with SSTs (R. [[#Zhang--2019|]] [[#Zhang--2019|]] [[#Zhang--2019|Zhang et al., 2019]] ), it is generally accepted that AMOC forcing of SST variability exists alongside stochastic wind forcing and external forcing by aerosols ( [[#Bellomo--2018|Bellomo et al., 2018]] ; [[#Haustein--2019|Haustein et al., 2019]] ; [[#O’Reilly--2019|O’Reilly et al., 2019]] ; [[#Wills--2019|Wills et al., 2019]] ).

The SROCC ( [[#Collins--2019|Collins et al., 2019]] ) assessed that in situ observations (2004–2017) and sea surface temperature reconstructions indicate that AMOC has weakened relative to 1850–1900 ( ''medium confidence'' ). However, SROCC also assessed that there is insufficient data to quantify the magnitude of the weakening, or to properly attribute it to anthropogenic forcing, due to the limited length of the observational record. Here, this assessment is adjusted to ''low confidence'' in the weakening (as also discussed in Sections 2.3.3.4.1 and 3.5.4.1). The CMIP5 multi-model mean showed no 20th century trend in AMOC ( [[#Cheng--2013|Cheng et al., 2013]] ). The CMIP6 multi-model mean slightly opposes the reconstructed decline due to a strong increase in the 1940–1985 period ( [[#Menary--2020b|Menary et al., 2020b]] ; [[#Weijer--2020|Weijer et al., 2020]] ), thought to be in response to aerosol forcing ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.4.1|Section 3.5.4.1]] ), followed by a smaller decline since the 1990s. Also, agreement between different proxy-based reconstructions is weak in many details ( [[#Moffa-Sánchez--2019|Moffa-Sánchez et al., 2019]] ) and questions can be raised regarding various proxies used in reconstructions ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4.1|Section 2.3.3.4.1]] ). For instance, SST-based proxies can be influenced by atmospheric and other processes acting on different time scales ( [[#Moffa-Sánchez--2019|Moffa-Sánchez et al., 2019]] ; [[#Jackson--2020|Jackson and Wood, 2020]] ). In addition, many proxies are indirect and based on AMOC-related processes assumed to be similar to those found in models, such as the link between AMOC and Labrador Sea convection, which has been questioned recently (see above). In addition, the subpolar gyre from which many AMOC proxies are taken may vary independently of AMOC, with similar patterns in SST and OHC driven by wind variability ( [[#Williams--2014|Williams et al., 2014]] ; [[#Piecuch--2017|Piecuch et al., 2017]] ). Finally, a new dynamic reconstruction of the Atlantic inflow to the Nordic Seas suggests no slowdown over the past 70 to 100 years ( [[#Rossby--2020|Rossby et al., 2020]] ), in contrast to a new compilation of proxy reconstructions which suggests that AMOC is presently in its weakest state in the last millennium ( [[#Caesar--2021|Caesar et al., 2021]] ), reinforcing the evidence that motivated the previous SROCC assessment. [[IPCC:Wg1:Chapter:Chapter-3#3.5.4.1|Section 3.5.4.1]] also questions the veracity of the models’ forced AMOC response during the 20th century. Given the large discrepancy between modelled and reconstructed AMOC in the 20th century, and the uncertainty over the realism of the 20th century modelled AMOC response ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.4.1|Section 3.5.4.1]] ), we have ''low confidence'' in both.

The strength of AMOC has been measured directly since 2004 using the RAPID Array ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4.1|Section 2.3.3.4.1]] ; [[#Smeed--2018|Smeed et al., 2018]] ). RAPID-based estimates show a large amount of variability compared to CMIP models ( [[#Roberts--2014|Roberts et al., 2014]] ). Observed changes since 2004 are too short for the evaluation of a long-term trend given the decadal scale internal variability ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4.1|Section 2.3.3.4.1]] ). Nevertheless, [[#Smeed--2018|Smeed et al. (2018)]] argue that, between 2007 and 2011, AMOC shifted to a state of reduced overturning – decreasing from 18.8 Sv between 2004 and 2008 to 16.1 Sv after 2008. A shift in AMOC strength of this magnitude is not captured by CMIP5 and CMIP6 models, which generally underestimate interannual to decadal AMOC variability ( [[IPCC:Wg1:Chapter:Chapter-3#3.5.4.1|Section 3.5.4.1]] ). Additional evidence since SROCC also raises the inconsistency between the RAPID weakening in the 3000–5000 m depth range and the relative constancy of deep overflows from the Arctic ( [[#Østerhus--2019|Østerhus et al., 2019]] ), implying that the recent decrease in AMOC at 26.5°N ( [[#Smeed--2018|Smeed et al., 2018]] ) is not caused by overflow weakening or reduced overturning in the Nordic Seas, although the weakening occurred almost exclusively in the 3000–5000 m depth range associated with a reduction of Lower NADW ( [[#9.2.2.3|Section 9.2.2.3]] ). It is unclear what causes a weakening of the deepest limb of AMOC at 26.5°N, if the main sources for this flow farther north remain constant. Various estimates of AMOC and associated heat transport suggest an increase since the 1940s with a subsequent decrease since the 1990s ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4.1|Section 2.3.3.4.1]] ), supported by ocean reanalysis ( [[#Jackson--2019|Jackson et al., 2019]] ), forced ocean model simulations ( [[#Robson--2012|Robson et al., 2012]] ; [[#Danabasoglu--2016|Danabasoglu et al., 2016]] ) and CMIP6 simulations ( [[#Menary--2020a|Menary et al., 2020a]] ). This suggests that the observed AMOC-shift between 2007 and 2011 may be part of a longer-term decrease ( ''medium confidence'' ), which has been attributed to be part of multiannual variability ( [[#Rhein--2019|Rhein et al., 2019]] ).

The SROCC ( [[#Collins--2019|Collins et al., 2019]] ) found that AMOC will ''very likely'' weaken over the 21st century. In CMIP6 projections, the modelled decline starting in the 1990s continues in all future projections, almost independent of the forcing scenario until about 2060, after which low-emissions scenarios show stabilization, while high-emissions scenarios continue to exhibit AMOC decline (Figure 9.10; [[#Menary--2020b|Menary et al., 2020b]] ; [[#Weijer--2020|Weijer et al., 2020]] ). Despite differences in overall AMOC strength, location and latitude of deep convection, sea ice and SST bias and representation of deep overflows, the model projections are qualitatively similar. This agreement suggests that AMOC decline may be governed by large-scale constraints independent of the details of the models. In theoretical models of the thermohaline circulation, the circulation strength is proportional to a density or pressure difference between the subpolar North Atlantic and subtropical South Atlantic ( [[#Kuhlbrodt--2007|Kuhlbrodt et al., 2007]] ; [[#Weijer--2019|Weijer et al., 2019]] ). In all models, the north-south pressure gradient decreases in the 21st century, as subpolar waters warm faster than subtropical waters, and an enhanced hydrological cycle drives freshening at subpolar latitudes, while subtropical latitudes feature more evaporation and salinification ( [[#9.2.1|Section 9.2.1]] ). As a result, surface waters at subpolar latitudes become more buoyant and more stable, so that deep water formation driving the AMOC declines ( [[#9.2.1.3|Section 9.2.1.3]] ). Projected AMOC decline by 2100 ranges from 24 [4 to 46] % in SSP1-2.6 to 39 [17–55] % in SSP5-8.5 ( ''medium confidence'' ) ( [[IPCC:Wg1:Chapter:Chapter-4#4.3.2.3|Section 4.3.2.3]] ). Note that these ranges are based on ensemble means of individual models, largely smoothing out internal variability. If single realizations are considered, the ranges become wider, especially by lowering the low end of the range ( [[IPCC:Wg1:Chapter:Chapter-4#4.3.2.3|Section 4.3.2.3]] ). In summary, it is ''very likely'' that AMOC will decline in the 21st century, but there is ''low confidence'' in the model’s projected timing and magnitude. In addition, freshwater from the melting of the Greenland Ice Sheet (Sections 9.4.1.3 and 9.4.1.4) could further enhance the future weakening of AMOC in the 21st century ( [[#Collins--2019|Collins et al., 2019]] ; [[#Golledge--2019|Golledge et al., 2019]] ).

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[[File:59bfd54f25dac32fb13e065da36f3646 IPCC_AR6_WGI_Figure_9_10.png]]

'''Figure 9.10''' '''|''' '''Atlantic Meridional Overturning Circulation (AMOC) strength in simulations and sensitivity to resolution and forcing.''' '''(Top left)''' AMOC magnitude (units: Sverdrup (Sv) = 10 <sup>9</sup> kg <sup></sup> s <sup>–1</sup> ) in Paleoclimate Modelling Intercomparison Project (PMIP) experiments. '''(Top right)''' Time series of AMOC from Coupled Model Intercomparison Project Phase 5 and 6 (CMIP5 and CMIP6) based on ( [[#Menary--2020b|Menary et al., 2020b]] ). '''(Bottom left)''' Percent change in AMOC strength per year at different resolutions over the 1950–2050 period with colours for model families ( [[#Roberts--2020|Roberts et al., 2020]] ). '''(Bottom right)''' A compilation of percentage changes in the simulated AMOC after applying an additional freshwater flux in the subpolar North Atlantic at the surface for a limited time ( [[#de%20Vries--2005|de Vries and Weber, 2005]] ; [[#Stouffer--2006|Stouffer et al., 2006]] ; [[#Yin--2007|Yin and Stouffer, 2007]] ; [[#Jackson--2013|Jackson, 2013]] ; [[#Liu--2013|Liu and Liu, 2013]] ; [[#Jackson--2018|Jackson and Wood, 2018]] ; [[#Haskins--2019|Haskins et al., 2019]] ). Symbols indicate whether the AMOC recovers within 200 years (circles), is starting to recover (upwards arrow), or does not recover within 200 years (downwards arrow). Symbol size indicates rate of freshwater input. Further details on data sources and processing are available in the chapter data table (Table 9.SM.9).

Both AR5 ( [[#Collins--2013|Collins et al., 2013]] ) and SROCC ( [[#Collins--2019|Collins et al., 2019]] ) assessed that an abrupt collapse of AMOC before 2100 was ''very unlikely'' , but SROCC added that, by 2300, an AMOC collapse was ''as likely asnot'' for high-emissions scenarios. The SROCC also assessed that model bias may considerably affect the sensitivity of the modelled AMOC to freshwater forcing. Tuning towards stability and model biases ( [[#Valdes--2011|Valdes, 2011]] ; [[#Liu--2017|Liu et al., 2017]] ; [[#Mecking--2017|Mecking et al., 2017]] ; [[#Weijer--2019|Weijer et al., 2019]] ) provides CMIP models a tendency toward unrealistic stability ( ''medium confidence'' ). By correcting for existing salinity biases, [[#Liu--2017|Liu et al. (2017)]] demonstrated that AMOC behaviour may change dramatically on centennial to millennial time scales, and that the probability of a collapsed state increases. None of the CMIP6 models features an abrupt AMOC collapse in the 21st century, but they neglect meltwater release from the Greenland Ice Sheet. Also, a recent process study reveals that a collapse of AMOC can be induced, even by small-amplitude changes in freshwater forcing ( [[#Lohmann--2021|Lohmann and Ditlevsen, 2021]] ). As a result, we change the assessment of an abrupt collapse before 2100 to ''medium confidence'' that it will not occur ''.''

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<span id="southern-ocean"></span>
==== 9.2.3.2 Southern Ocean ====

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The changing Southern Ocean circulation system exerts a strong influence on the global climate by modulating: (i) global OHC ( [[#9.2.2.1|Section 9.2.2.1]] ); (ii) global ocean anthropogenic carbon uptake (Cross-chapter Box 5.3); global ocean overturning circulation ( [[#9.2.3.1|Section 9.2.3.1]] ); (iii) climate sensitivity ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4|Section 7.4.4]] and Cross-chapter Box 5.3); (iv) sea level through basal melt of ice shelves (9.4.2); and (v) Southern Hemisphere sea ice cover ( [[#9.3.2|Section 9.3.2]] ).

The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) had ''low confidence'' in all CMIP5-based model projections due to their inability to explicitly resolve eddy processes, and their inability to properly consider future meltwater change from the Antarctic Ice Sheet. These limitations of climate models to represent the Southern Ocean persist due to most CMIP6 models still using parameterized mesoscale eddy processes, which are limited in projecting the future response of the horizontal and vertical circulation under climate warming, and also because of the continued absence of active ice-shelf and ice-sheet coupling in the CMIP6 model suite, therefore ignoring basal meltwater and calving feedback on the circulation ( [[#Meredith--2019|Meredith et al., 2019]] ). In addition, two important limitations of CMIP6 models of the Southern Ocean involve processes that were not assessed in SROCC. First, the poor representation of dense overflows causes most of the Antarctic Bottom Water (AABW) to be formed by spurious open ocean convection rather than by dense overflows from the Antarctic continental shelves that feed the lower overturning cell ( [[#Snow--2015|Snow et al., 2015]] ; [[#Dufour--2017|Dufour et al., 2017]] ; [[#Heuzé--2021|Heuzé, 2021]] ). Second, Antarctic continental shelf waters are poorly simulated because potentially important controlling mechanisms tend to be too small and transient to observe and resolve in CMIP ocean models. These small processes include: the heterogeneity of observed sub-ice-shelf melt with warm water driving narrow basal channels that cut underneath the ice ( [[#Drews--2015|Drews, 2015]] ; [[#Alley--2016|Alley et al., 2016]] ; [[#Marsh--2016|Marsh et al., 2016]] ; [[#Milillo--2019|Milillo et al., 2019]] ); eddies and tides ( [[#Stewart--2018|Stewart et al., 2018]] ; [[#Jourdain--2019|Jourdain et al., 2019]] ; [[#Hausmann--2020|Hausmann et al., 2020]] ), which can drive Circumpolar Deep Water (CDW) onto the continental shelves or dynamically increase melting ( [[#9.2.3.6|Section 9.2.3.6]] ); and feedback mechanisms between ocean, atmosphere and cryosphere that can weaken or amplify initial perturbations ( [[#Donat-Magnin--2017|Donat-Magnin et al., 2017]] ; [[#Spence--2017|Spence et al., 2017]] ; [[#Turner--2017|Turner et al., 2017]] ; [[#Silvano--2018|Silvano et al., 2018]] ; [[#Webber--2019|Webber et al., 2019]] ; [[#Hazel--2020|Hazel and Stewart, 2020]] ). In addition, the Southern Ocean in CMIP5 and CMIP6 models exhibit surface temperature biases ( [[#9.2.1.1|Section 9.2.1.1]] ), which have been linked in CMIP5 models to errors in atmospheric model cloud-related shortwave radiation ( [[#Hyder--2018|Hyder et al., 2018]] ) and are somewhat improved in High Resolution Model Intercomparison Project (HighResMIP) models (Figure 9.3). In summary, there is ''high confidence'' that future change in the subpolar Southern Ocean region, including sea ice cover and ocean temperature change on Antarctic continental shelves, depends on feedback mechanisms involving the ocean, atmosphere and cryosphere that are poorly understood and not represented in the current generation of climate models. This results in large uncertainty and ''low confidence'' in the future sea ice cover ( [[#9.3.2|Section 9.3.2]] ) and in ocean temperature change on the Antarctic continental shelf ( [[#9.4.2.3|Section 9.4.2.3]] ).

Despite these challenges, the CMIP6 ensemble does represent the main Southern Ocean circulation characteristics: the simulated Antarctic Circumpolar Current (ACC) transport is generally lower than observation-based values but consistent when considering ensemble spread, and the inter-model spread in ACC transport has greatly reduced from previous generations of climate models from CMIP3 to CMIP6 ( [[#Beadling--2019|Beadling et al., 2019]] , 2020). The structure (but not the magnitude) of the two-cell zonally averaged overturning is captured by most CMIP6 models ( [[#Russell--2018|Russell et al., 2018]] ; [[#Beadling--2019|Beadling et al., 2019]] ). In addition, while issues remain, CMIP6 climate models show clear improvements in their representation of AABW compared to CMIP5: several models correctly represent or parameterize Antarctic shelf processes, fewer models exhibit Southern Ocean deep convection, bottom density biases are reduced, and abyssal overturning is more realistic ( [[#Heuzé--2021|Heuzé, 2021]] ). In terms of atmospheric wind forcing, CMIP6 models show an improvement compared to CMIP5 models, with an overall reduction in the equatorward bias of the annual mean westerly jet from 1.9° in CMIP5 to 0.4° in CMIP6, but in contrast, they show no such overall improvements for their representation of the Amundsen Sea Low ( [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ; [[#Lyu--2020a|Lyu et al., 2020a]] ), which can be critical in driving variability of water masses on the Antarctic continental shelf in west Antarctica, the Weddell Sea or the Ross Sea ( [[#Holland--2019|Holland et al., 2019]] ; [[#Silvano--2020|Silvano et al., 2020]] ).

The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) established that, while trends in the atmospheric forcing of the Southern Ocean have been dominated by a strengthening of the Southern Hemisphere westerly winds in recent decades, there is ''medium confidence'' that ACC transport is weakly sensitive to changes in winds. It also reported that, instead of increasing the mean ACC transport, additional energy input associated with increased wind stress cascades into the eddy field ( ''medium confidence'' ). In contrast with the AR5 assessment ( [[#Rhein--2013|Rhein et al., 2013]] ), SROCC evaluated that it was ''unlikely'' that there has been a net southward migration of the mean ACC position over the past 20 years. There is no additional evidence to revisit SROCC assessment on wind sensitivity. However, new evidence does suggest that air–sea buoyancy forcing associated with idealized 4×CO <sub>2</sub> forcing leads to an increase in ACC transport ( ''limited evidence'' ) ( [[#Shi--2020|Shi et al., 2020]] ). The SROCC noted that, if the general strengthening in westerly winds is sustained, then it is ''very likely'' that the eddy field will continue to increase in intensity, and it is ''likely'' that the mean position and strength of the ACC will remain only weakly sensitive to winds. In the future, the strength of the Southern Hemisphere westerly wind jet results from a competition between decrease due to ozone hole recovery and increase due to increased radiative forcing ( [[IPCC:Wg1:Chapter:Chapter-4#4.3.3.1|Section 4.3.3.1]] ). This competition results in an increased atmospheric jet by 2100 compared to present day under SSP2-4.5, SSP3-7.0, and SSP5-8.5, but a decreased jet by 2100 under SSP1-2.6 ( [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ). There is little inter-model spread in the CMIP6 future response of the atmospheric westerly jet, providing ''high confidence'' in this assessment (in contrast, CMIP6 models show no consistency in their future projection of easterly wind change along the Antarctic continental shelf break; [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ). Paleo-oceanographic evidence suggests that ACC flow through Drake Passage was consistently stronger during warm intervals of the past (both during interstadials and interglacials), but with relatively little change and no consensus on the sign of change in other regions ( [[#Lamy--2015|Lamy et al., 2015]] ; [[#Toyos--2020|Toyos et al., 2020]] ). In summary, additional evidence since SROCC confirms that there is ''medium confidence'' that the ACC has been weakly sensitive to Southern Hemisphere atmospheric jet increase in the past decades. New evidence since SROCC suggests that there is ''high confidence'' that the Southern Hemisphere atmospheric jet will increase in the 21st century for all scenarios (except for SSP1-1.9 and SSP1-2.6; [[IPCC:Wg1:Chapter:Chapter-4#4.3.3.1|Section 4.3.3.1]] ) with a greater increase for larger radiative forcing. An increase in westerly winds will ''very likely'' force an increase of the eddy field in the ACC, and while there is ''medium confidence'' that the ACC is weakly sensitive to wind change, new advances since SROCC provide ''limited evidence'' that the ACC transport will nevertheless increase in response to wind and buoyancy fluxes.

For the upper cell overturning circulation, SROCC concluded that: its transport has experienced significant inter-decadal variability in response to wind forcing since the 1990s; and there is ''low confidence'' in the assessments of a long-term increase in upper-ocean overturning. Consistent with SROCC, the importance of eddy processes and winds in driving long-term change and variability have been reinforced, with a potential fast wind response partially counteracted by a slower eddy response ( [[#Doddridge--2019|Doddridge et al., 2019]] ; [[#Waugh--2019|Waugh et al., 2019]] ; [[#Stewart--2020|Stewart et al., 2020]] ). Eddy parametrizations affect the strength of overturning, its sensitivity to winds and the ACC transport ( [[#Mak--2017|Mak et al., 2017]] ). Even in eddy-resolving simulations, sub-gridscale dissipation affects the overturning and ACC ( [[#Pearson--2017|Pearson et al., 2017]] ). In addition, there has been progress in understanding the importance of Antarctic Ice Shelf meltwater and sea ice, in driving the observed changes in the near surface and in the upper overturning cell over the past decades, on top of changes induced by winds and eddies ( [[#Bronselaer--2020|Bronselaer et al., 2020]] ; [[#Haumann--2020|Haumann et al., 2020]] ; [[#Jeong--2020|Jeong et al., 2020]] ; [[#Rye--2020|Rye et al., 2020]] ). In particular, increased stratification caused by increased freshwater flux to the surface ocean ( [[#9.2.1.3|Section 9.2.1.3]] ) can cause a shoaling and warming of the CDW layer, and create a positive feedback, enhancing basal melt of the Antarctic Ice Sheet ( [[#9.4.2.1|Section 9.4.2.1]] ; [[#Bronselaer--2018|Bronselaer et al., 2018]] ; [[#Golledge--2019|Golledge et al., 2019]] ; [[#Schloesser--2019|Schloesser et al., 2019]] ; [[#Sadai--2020|Sadai et al., 2020]] ). There is ''medium confidence'' in the existence of this feedback mechanism but ''low agreement'' on the magnitude of the feedback. The SROCC reported that CMIP5 models project that the overall transport of upper-ocean overturning cell will increase by up to 20% in the 21st century, and no new studies alter that assessment.

For the lower cell overturning circulation, SROCC assessed that a slowdown of its transport is consistent with the observed decrease in volume ( ''medium confidence'' ) of AABW in the global ocean ( [[#9.2.2.3|Section 9.2.2.3]] ). Additional evidence since SROCC strengthens confidence that increased glacial meltwater flux will reduce the density of bottom waters during the 21st century. It will eventually reach a point where deep convection will be curtailed, and shelf water will become too buoyant to sink to the ocean interior, thereby slowing the lower cell overturning circulation ( [[#Bronselaer--2018|Bronselaer et al., 2018]] ; [[#Golledge--2019|Golledge et al., 2019]] ; [[#Lago--2019|Lago and England, 2019]] ; [[#Moorman--2020|Moorman et al., 2020]] ). While such changes are consistent with the observed freshening and decreased volume of the AABW layer reported in SROCC (as discussed in [[#9.2.2.3|Section 9.2.2.3]] ), new observation-based studies have highlighted how the lower cell overturning can episodically increase as a response to climate anomalies, temporally counteracting the tendency for melt to reduce AABW formation ( [[#Abrahamsen--2019|Abrahamsen et al., 2019]] ; [[#Castagno--2019|Castagno et al., 2019]] ; [[#Gordon--2020|Gordon et al., 2020]] ; [[#Silvano--2020|Silvano et al., 2020]] ). In addition, while the opening of open ocean polynyas can affect the lower cell on decadal to centennial time scales, there is ''limited evidence'' and ''low agreement'' in the role of open ocean polynyas in driving observed trends of the lower cell in the last decade ( [[#9.2.2.3|Section 9.2.2.3]] ). Based on CMIP5 models, SROCC reported with ''low confidence'' that formation and export of AABW associated with the lower overturning cell will decrease in the 21st century, and there is no new evidence to revisit that assessment from climate models. However, additional paleo evidence from marine sediments suggests that AABW formation/ventilation was vulnerable to freshwater fluxes during past interglacials ( [[#Hayes--2014|Hayes et al., 2014]] ; [[#Huang--2020|Huang et al., 2020]] ; [[#Turney--2020|Turney et al., 2020]] ) and that AABW formation was strongly reduced ( [[#Skinner--2010|Skinner et al., 2010]] ; [[#Gottschalk--2016|Gottschalk et al., 2016]] ; [[#Jaccard--2016|Jaccard et al., 2016]] ) or possibly totally curtailed ( [[#Huang--2020|Huang et al., 2020]] ) during the Last Glacial Maximum (LGM) and transient cold intervals of marine isotope stages 2 and 3 (MIS2 and MIS3). Specifically, sedimentary reconstructions show a transient reduction in AABW ventilation in the Atlantic sector of the Southern Ocean during MIS5e, which is assessed to have been warmer than modern climate ( [[#Thomas--2020|Thomas et al., 2020]] ). However, long multi-centennial or millennial model runs under higher-than-pre-industrial CO <sub>2</sub> concentrations show that, after 500–1000 years, ventilation in the Southern Ocean resumes, and possibly overshoots with enhanced convection in the Weddell and Ross seas, leading to enhanced bottom water ventilation globally ( [[#Yamamoto--2015|Yamamoto et al., 2015]] ; [[#Frölicher--2020|Frölicher et al., 2020]] ). AABW ventilation increased at the onset of the last deglacial transition, promoting the release of previously sequestered CO <sub>2</sub> to the atmosphere on centennial to millennial time scales ( [[#Bauska--2016|Bauska et al., 2016]] ; [[#Jaccard--2016|Jaccard et al., 2016]] ; [[#Rae--2018|Rae et al., 2018]] ), concomitant with a southward shift of the Southern Hemisphere westerly wind belt ( [[#Denton--2010|Denton et al., 2010]] ; [[#Jaccard--2016|Jaccard et al., 2016]] ) and reduced sea ice cover ( [[#Ferrari--2014|Ferrari et al., 2014]] ; [[#Stein--2020|Stein et al., 2020]] ). In summary, the combination of observational, numerical and paleoclimate evidence provides us with ''medium confidence'' that the lower cell will continue decreasing in the 21st century as a result of increased basal melt from the Antarctic Ice Sheet.

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<span id="tropical-oceans"></span>
==== 9.2.3.3 Tropical Oceans ====

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The tropics are a tightly coupled ocean-atmosphere system with tightly interconnected basins ( [[#Cai--2019|Cai et al., 2019]] ). The zonal atmospheric Walker Circulation and the Indonesian Throughflow (Figure 9.11) are key connections between the Pacific and Indian oceans, and variations in the Walker and Hadley Circulations are tightly linked to the tropical Pacific SST and currents. The tropics have a profound influence on the climate system through the multiple modes of variability they host, which have widespread global influence at seasonal to annual time scale (Annex IV).

The effect of tropical modes of variability on climate and their long-term changes are reviewed in detail in Annex IV, while changes to the tropical ocean are assessed throughout the report and briefly summarized here. [[IPCC:Wg1:Chapter:Chapter-2#2.4|Section 2.4]] concludes that a sustained shift beyond multi-centennial variability has not been observed for El Niño–Southern Oscillation (ENSO) ( ''medium confidence'' ) and that there is ''limited evidence'' and ''limited agreement'' about the long-term behaviour of other tropical modes. [[IPCC:Wg1:Chapter:Chapter-3#3.7|Section 3.7]] assesses with ''high confidence'' that human influence has not affected the principal tropical modes of interannual climate variability and their associated regional teleconnections beyond the range of internal variability. [[IPCC:Wg1:Chapter:Chapter-4#4.3.3.2|Section 4.3.3.2]] assesses with ''medium confidence'' that there is no consensus from models for a systematic change in the amplitude of ENSO sea surface temperature variability over the 21st century. The related change in tropical SSTs is covered in [[#9.2.1.1|Section 9.2.1.1]] . The projected changes in SST have implications for marine heat wave characteristics, which are assessed in Box 9.2. SST changes in the tropics are related to changes in the atmospheric circulation, including surface equatorial easterly trade winds and Walker Circulation ( [[IPCC:Wg1:Chapter:Chapter-4#4.5.3.2|Section 4.5.3.2]] ), and the weakening Indonesian Throughflow and strengthening Agulhas Extension and leakage ( [[#9.2.3.4|Section 9.2.3.4]] ). Weakening trade winds under climate change ( [[#Vecchi--2007|Vecchi and Soden, 2007]] ) will tend to decrease upwelling, along isopycnals in the eastern Pacific and diapycnal upwelling in the central Pacific, and thus the meridional temperature gradients that drive tropical instability waves ( [[#Terada--2020|Terada et al., 2020]] ), along with a weakening, flattening and shoaling of the tropical thermocline and equatorial undercurrent ( [[#Luo--2011|Luo and Rothstein, 2011]] ). A weak or absent equatorial undercurrent ( [[#Kuntz--2020|Kuntz and Schrag, 2020]] ) and a too-diffuse and incorrectly sloped tropical thermocline ( [[#Zhu--2020|Zhu et al., 2020]] ) remain issues in most CMIP6 models. In summary, while future changes in tropical modes of variability remain unclear, change in atmospheric and ocean circulation will drive continued change in tropical ocean temperature in the 21st century ( ''medium confidence'' ), with part of the region experiencing drastic marine heat wave conditions ( ''high confidence'' ).

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<span id="gyres-western-boundary-currents-and-inter-basin-exchanges"></span>
==== 9.2.3.4 Gyres, Western Boundary Currents and Inter-basin Exchanges ====

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The AR5 ( [[#Rhein--2013|Rhein et al., 2013]] ) assessed with ''medium'' to ''high confidence'' that the North Pacific subpolar gyre, the South Pacific subtropical gyre, and the subtropical cells have intensified. They also reported that the North Pacific subtropical gyre had expanded since the 1990s, and that, overall, the changes in gyre systems were ''likely'' predominantly due to interannual-to-decadal variability. The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) complemented the AR5 assessment by reporting that the polar Beaufort Gyre in the Arctic expanded to the north-west between 2003 and 2014, contemporaneous with changes in its freshwater accumulation and alterations in wind forcing. Consistent with the reported change over the gyres, both AR5 and SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ; [[#Collins--2019|Collins et al., 2019]] ) reported that western boundary currents (WBCs) have intensified (Figure 9.11), and expanded poleward, except for the Gulf Stream and the Kuroshio. [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4|Section 2.3.3.4]] provides an overall assessment of gyres and WBCs, including an assessment of change from paleoclimate archives. [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4|Section 2.3.3.4]] assesses that, while WBC strength is highly variable at multi-decadal scale ( ''high confidence'' ), WBCs and subtropical gyres have shifted poleward since 1993 ( ''medium confidence'' ), at a rate on the order of 0.04–0.1 degree per decade during 1993–2018. Figure 9.11 shows that CMIP5 and CMIP6 models agree in projecting a weaker Gulf Stream and Gulf Stream Extension, while the Kuroshio changes less ( [[#Sen%20Gupta--2016|Sen Gupta et al., 2016]] ).

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[[File:84e0a170685aa8fde4ea7956a4aaed8e IPCC_AR6_WGI_Figure_9_11.png]]

'''Figure 9.11''' '''|''' '''Simulated barotropic streamfunction, surface speed and major current transport in Coupled Model Intercomparison Project Phase 5 and 6 (CMIP5 and CMIP6).''' '''(a)''' Mean barotropic streamfunction (unit: 10 <sup>9</sup> kg <sup></sup> s <sup>–1</sup> ; 1995–2014) and projected barotropic streamfunction change (10 <sup>9</sup> kg <sup></sup> s <sup>–1</sup> ; 2018–2100 vs 1995–2014) under '''(b)''' SSP5-8.5. '''(d)''' Mean surface (0–100 m) speed (m s <sup>–1</sup> ) and projected surface speed change (m s <sup>–1</sup> , 2081–2100) versus 1995–2014 under '''(e)''' SSP5-8.5. '''(c, f)''' Median and likely range of 1995–2014 and 2081–2100 transport of three currents with the largest transport change and four with the largest fractional change ( [[#Sen%20Gupta--2016|Sen Gupta et al., 2016]] ). '''(c)''' Deep currents: Agulhas Extension (ACx), Gulf Stream (GS), Gulf Stream Extension (GSx), Tasman Leakage (TASL), East Australia Current Extension (EACx), Indonesian Throughflow (ITF), and Brazil Current (BC). '''(f)''' Shallow currents: as for deep but with New Guinea Current (NGC), and without ACx. 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).

Although the observed wind stress curl shows systematic poleward shift in each basin as a result of anthropogenic warming ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1.4|Section 2.3.1.4]] ; [[#Chen--2012|Chen and Wu, 2012]] ; [[#Wu--2012|Wu et al., 2012]] ; [[#Zhai--2014|Zhai et al., 2014]] ), which has caused a systematic shift of the WBCs and subtropical gyres since 1993 ( [[#Wu--2012|Wu et al., 2012]] ; [[#Yang--2016|Yang et al., 2016]] , 2020), the response of current strength is more complex and inconsistent across regions ( [[#Sloyan--2015|Sloyan and O’Kane, 2015]] ; [[#Wang--2016|Y.-L. Wang et al., 2016]] ; [[#Elipot--2018|Elipot and Beal, 2018]] ; [[#McCarthy--2018|McCarthy et al., 2018]] ; [[#Wang--2018|Wang and Wu, 2018]] ; [[#Dong--2019|Dong et al., 2019]] ). The strength of WBCs and gyres exhibit inconsistent responses because they are dependent on wind stress forcing and because multi-scale interaction and air–sea interaction have an important role in their long-term trends and variability ( [[#Zhang--2020|Zhang et al., 2020]] ). Observed changes in gyre circulation are dominated by interannual and decadal modes of variability globally ( [[#Qiu--2012|Qiu and Chen, 2012]] ; [[#Melzer--2017|Melzer and Subrahmanyam, 2017]] ; [[#McCarthy--2018|McCarthy et al., 2018]] ; [[#Hu--2020|Hu et al., 2020]] ). The North Atlantic subpolar gyre is strongly modulated by variability associated with the NAO and AMV (Annex IV; [[#Robson--2016|Robson et al., 2016]] ). Subpolar gyre systems can change abruptly due to a positive feedback between convective mixing and salinity transport ( [[#Born--2013|Born et al., 2013]] , 2016) and air–sea interaction ( [[#Moffa-Sánchez--2014|Moffa-Sánchez et al., 2014]] ; [[#Moreno-Chamarro--2017|Moreno-Chamarro et al., 2017]] ) within the gyre. In the Arctic, both the Beaufort gyre and mesoscale eddies strengthened between 2003 and 2014 ( [[#Armitage--2017|Armitage et al., 2017]] ), which might be partly due to increased wind stress ( [[#Oldenburg--2018|Oldenburg et al., 2018]] ) or reduced sea ice thickness and changes in sea ice pack morphology ( [[#van%20der%20Linden--2019|van der Linden et al., 2019]] ). Presently, there is ''limited evidence'' in attributing causality to these changes for any of the proposed mechanisms. In the North Pacific, there has been an increasing trend in the Alaska Gyre from 1993 to 2017 ( [[#Cummins--2018|Cummins and Masson, 2018]] ), which might be attributed to Pacific Decadal Oscillation ( ''low confidence'' ) ( [[#Hristova--2019|Hristova et al., 2019]] ). In the Southern Ocean, ''limited evidence'' indicates that the subpolar gyres respond to Southern Hemisphere atmospheric modes of variability at interannual time scale ( [[#Armitage--2018|Armitage et al., 2018]] ; [[#Dotto--2018|Dotto et al., 2018]] ).

All climate models reproduce WBCs and gyres, but eddy-present or eddy-rich models (roughly 10–25 km and about 10 km resolution, respectively) represent these currents more realistically than eddy-parameterized models ( ''very'' ''high confidence'' ) ( [[#Small--2014|Small et al., 2014]] ; [[#Griffies--2015|Griffies et al., 2015]] ; [[#Chassignet--2017|Chassignet et al., 2017]] , 2020; [[#Hewitt--2017|Hewitt et al., 2017]] , 2020; [[#Roberts--2018|Roberts et al., 2018]] ). Compared to observations or to eddy-present and eddy-rich models, the eddy-parameterized models from CMIP5 and CMIP6 simulate weaker and wider WBCs, as well as less realistic locations of subtropical and subpolar gyre boundaries (Figure 9.11). Increased resolution admits mesoscale eddies, and also improves simulation of the strength and position of WBCs such as the Kuroshio Current, Gulf Stream, and East Australian Current ( ''very high confidence'' ) ( [[#Sasaki--2004|Sasaki et al., 2004]] ; [[#Chassignet--2008|Chassignet and Marshall, 2008]] ; [[#Delworth--2012|Delworth et al., 2012]] ; [[#Yu--2012|Yu et al., 2012]] ; [[#Small--2014|Small et al., 2014]] ; [[#Haarsma--2016|Haarsma et al., 2016]] ; [[#Chassignet--2017|Chassignet et al., 2017]] , 2020; [[#Hewitt--2020|Hewitt et al., 2020]] ). Improved boundary current location relates to improved recirculation regions ( [[#Jayne--2009|Jayne et al., 2009]] ), mean path and variability, and existence of multiple stable paths ( [[#Qiu--2005|Qiu et al., 2005]] ; [[#Delman--2015|Delman et al., 2015]] ), air–sea fluxes ( [[#Small--2014|Small et al., 2014]] ), and related coastal weather patterns ( [[#Kaspi--2011|Kaspi and Schneider, 2011]] ). The wind-current feedback, implemented by considering relative velocity of currents and wind, realistically dampens mesoscale eddies and WBCs, through mesoscale air–sea interaction ( [[#Ma--2016|Ma et al., 2016]] ; [[#Renault--2016|Renault et al., 2016]] , 2019), even though sub-mesoscale wind-current damping feedback is missing in these models ( ''medium confidence'' ) (Z. [[#Zhang--2016|]] [[#Zhang--2016|]] [[#Zhang--2016|Zhang et al., 2016]] ). As eddies potentially play a role in determining the strength of gyre circulations and their low-frequency variability ( [[#Fox-Kemper--2004|Fox-Kemper and Pedlosky, 2004]] ; [[#Berloff--2007|Berloff et al., 2007]] ), it is expected that eddy-present and eddy-rich models will differ in their decadal variability and sensitivity to changes in the wind stress of gyres from eddy-parameterized models ( ''medium confidence'' ). Nonetheless, important aspects of gyre strength depend primarily on forcing and not resolution, allowing long-term changes in gyre strength to be investigated with low-resolution climate models ( [[#Hughes--2001|Hughes and de Cuevas, 2001]] ; [[#Yeager--2015|Yeager, 2015]] ).

Under future scenarios RCP4.5 and RCP8.5, AR5 ( [[#Collins--2013|Collins et al., 2013]] ) assessed an intensification and poleward extension of the southern Hemisphere subtropical gyres in the 21st century. New evidence since AR5 further reinforces their conclusions, which are now extended to all subtropical gyre systems in the Northern and Southern hemispheres ( [[#Yang--2016|Yang et al., 2016]] , 2020). CMIP6 models project changes in WBCs that are consistent with projected changes in the surface winds. Under strong radiative forcing, in scenario SSP5-8.5, CMIP6 models project that the East Australian Current Extension, Agulhas Current Extension and Brazil Current will intensify in the 21st century, while the Gulf Stream will weaken (Figure 9.11). Although CMIP5/CMIP6 are limited in resolution, ''medium confidence'' is given to changes in WBCs due to consistency across generations of climate models, including CMIP6, despite changes in model structure, resolution and parametrizations.

The SROCC ( [[#Collins--2019|Collins et al., 2019]] ) concluded with ''high confidence'' that Indonesian Throughflow (ITF) transport from the Pacific Ocean to the Indian Ocean has increased in the past two decades as a result ( ''medium confidence'' ) of an unprecedented intensification of the equatorial Pacific trade wind system. [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4|Section 2.3.3.4]] assesses that there is ''high confidence'' that the increase in the ITF over the past two decades is linked to multi-decadal scale variability rather than a longer-term trend. Consistently, in the future, as winds change under increased radiative forcing, most models project a decline of the ITF on the centennial time scale (Figure 9.11). One of the clearest changes of ocean current transport simulated by climate models is a weakening of the Indonesian Throughflow, projected in CMIP5 simulations under RCP4.5 and RCP8.5 scenarios ( [[#Sen%20Gupta--2016|Sen Gupta et al., 2016]] ; [[#Stellema--2019|Stellema et al., 2019]] ), and in CMIP6 simulations under the SSP5-8.5 scenario ( ''high confidence'' , Figure 9.11).

The SROCC reports with ''high confidence'' that the Agulhas leakage from the Indian to the Atlantic Ocean has increased in the past two decades ( [[#Collins--2019|Collins et al., 2019]] ), and there is no additional evidence since then allowing this assessment to be revisited ( [[#Biastoch--2015|Biastoch et al., 2015]] ; [[#Loveday--2015|Loveday et al., 2015]] ; [[#Lübbecke--2015|Lübbecke et al., 2015]] ). There is ''low confidence'' in future projections of Agulhas leakage because most CMIP models cannot directly simulate it, due to coarse resolution. However, there is ''medium evidence'' that the strength of the Southern Hemisphere westerlies controls Agulhas leakage ( [[#Durgadoo--2013|Durgadoo et al., 2013]] ; [[#Biastoch--2015|Biastoch et al., 2015]] ; [[#Loveday--2015|Loveday et al., 2015]] ), and ''high confidence'' that the strength of the Southern Hemisphere westerlies will increase under increased radiative forcing, except in lower warming scenarios (SSP1-1.9, SSP1.2-6; [[IPCC:Wg1:Chapter:Chapter-4#4.3.3.1|Section 4.3.3.1]] ; [[#Bracegirdle--2020|Bracegirdle et al., 2020]] ). There is also evidence that increasing Agulhas leakage is consistent with observed change of the temperature and salinity structure in the Atlantic ocean, and with variability of the AMOC ( [[#9.2.3.1|Section 9.2.3.1]] ; [[#Biastoch--2015|Biastoch et al., 2015]] ). This range of indirect evidence provides ''medium confidence'' that the Agulhas leakage will increase in the 21st century, except for the strongest mitigation scenario (Figure 9.11).

The SROCC assessed that the annual Bering Strait volume transport from the Pacific to the Arctic Ocean increased from 2001–2014, consistent with an estimated increased northward heat transport of about 60% from 2001–2014, and an increased freshwater transport of 30 ± 20 km <sup>3</sup> yr <sup>–1</sup> from 1991 to 2015 ( [[#Meredith--2019|Meredith et al., 2019]] ). [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4|Section 2.3.3.4]] assesses that volume transport from the Pacific to the Arctic has increased since the 1990s from 0.8 Sv to 1.0 Sv over 1990–2015. Realistic representation of the Bering Strait transport in the current generation of climate models is challenging because the strait is narrow compared to the resolution of climate models ( [[#Clement%20Kinney--2014|Clement Kinney et al., 2014]] ; [[#Aksenov--2016|Aksenov et al., 2016]] ). For the Atlantic to Arctic transport, [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4|Section 2.3.3.4]] reports that the major branches of Atlantic Water inflow across the Greenland–Scotland Ridge have remained stable, with only the smaller pathway of Atlantic Water north of Iceland showing a strengthening trend during 1993–2018. [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.4|Section 2.3.3.4]] also assesses that the Arctic outflow remained stable from the mid-1990s to the mid-2010s. Future changes in these currents have not yet been studied in CMIP6 models.

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<span id="eastern-boundary-upwelling-systems"></span>
==== 9.2.3.5 Eastern Boundary Upwelling Systems ====

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Eastern boundary upwelling systems (EBUS) exist where trade winds draw cold and generally low-pH/low-oxygen waters upward. Coastal upwelling plays a key role in supplying the food chain with nutrients, hence the richness and productivity of EBUS ( [[#Bindoff--2019|Bindoff et al., 2019]] ). The SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessed with ''high confidence'' that three out of the four major EBUS have experienced large-scale wind intensification in the past 60 years (only the trend for the Canary Current is considered uncertain). However, it also emphasized that various processes can also modulate, or even reverse, wind trends locally ( [[#Bindoff--2019|Bindoff et al., 2019]] ). Here we revisit SROCC assessment ( [[#Bindoff--2019|Bindoff et al., 2019]] ) based on evidence showing ''low agreement'' between studies that have investigated trends over past decadess of upwelling-favourable winds ( [[#Varela--2015|Varela et al., 2015]] ). This ''low agreement'' has been related to differences in wind products, season of interest, and length of the considered time series ( [[#Varela--2015|Varela et al., 2015]] ). Based on this, we assess that only the California Current system has experienced large-scale upwelling-favorable wind intensification over the period 1982–2010, albeit with regional differences ( [[#García-Reyes--2010|García-Reyes and Largier, 2010]] ; [[#Seo--2012|Seo et al., 2012]] ). In the Benguela, Canary, and Humboldt systems, large-scale, upwelling-favourable wind trends are ambiguous, owing to ''low confidence'' in long-term in situ marine wind data ( [[#Cardone--1990|Cardone et al., 1990]] ; [[#Bakun--2010|Bakun et al., 2010]] ) and ''low agreement'' among available studies ( [[#Narayan--2010|Narayan et al., 2010]] ; [[#Sydeman--2014|Sydeman et al., 2014]] ; [[#Varela--2015|Varela et al., 2015]] ). Our assessment confirms SROCC assessment ( [[#Bindoff--2019|Bindoff et al., 2019]] ) in that high natural variability of EBUS and their inadequate representation by most climate models gives ''low confidence'' in attribution of observed changes, while anthropogenic changes are projected to emerge primarily in the second half of the 21st century ( ''limited evidence'' : one model and one study) ( [[#Brady--2017|Brady et al., 2017]] ).

Under increased radiative forcing, SROCC ( [[#Bindoff--2019|Bindoff et al., 2019]] ) assessed that climate models project, in the 21st century, a reduction of wind and upwelling intensity in EBUS at low latitudes, and enhancement at high latitudes, under scenario RCP8.5, with an overall reduction in either upwelling intensity or extension. It also highlighted that coastal warming and wind intensification may lead to variable countervailing responses to upwelling intensification at local scales. Despite differences among EBUS (D. [[#Wang--2015|]] [[#Wang--2015|Wang et al., 2015]] ), there is growing evidence since SROCC in this pattern of change. While it has long been hypothesized that, for upwelling winds, change is linked to air temperature contrast between ocean and land ( [[#Bakun--1990|Bakun, 1990]] ), this hypothesis has increasingly been challenged. Changes in sea level pressure and wind fields in EBUS appear to be primarily tied to those affecting subtropical highs ( [[#García-Reyes--2013|García-Reyes et al., 2013]] ). Poleward expansion of the Hadley cell ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1.4.1|Section 2.3.1.4.1]] ; [[#Staten--2018|Staten et al., 2018]] ) and the related poleward migration of subtropical highs ( [[#He--2017|He et al., 2017]] ; [[#Cherchi--2018|Cherchi et al., 2018]] ), produce robust patterns of changes of reduced upwelling at low latitude and enhanced upwelling at high latitude ( [[#Echevin--2012|Echevin et al., 2012]] ; [[#Belmadani--2014|Belmadani et al., 2014]] ; [[#Bettencourt--2015|Bettencourt et al., 2015]] ; [[#Rykaczewski--2015|Rykaczewski et al., 2015]] ; [[#Sousa--2017|Sousa et al., 2017]] ; [[#Lamont--2018|Lamont et al., 2018]] ; [[#Sylla--2019|Sylla et al., 2019]] ). These patterns are most apparent in summer in both hemispheres. Synoptic variability of upwelling winds, important to the functioning of upwelling ecosystems ( [[#García-Reyes--2014|García-]] [[#Reyes--2014|Reyes et al., 2014]] ), may also be affected by climate change ( [[#Aguirre--2019|Aguirre et al., 2019]] ). However, coarse resolution model projections of winds in upwelling regions may be more consistent than higher-resolution projections, as these regions are highly sensitive to resolution ( [[#Small--2015|Small et al., 2015]] ).

Projected future annual cumulative upwelling wind changes at most locations, and seasons remain within ±10–20% of present-day values in the 21st century, even in the context of high-end emissions scenarios (4×CO <sub>2</sub> or RCP8.5) ( ''medium confidence'' ). Changes due to wind stress curl and alongshore pressure gradients tend to agree with alongshore wind changes ( [[#Oerder--2015|Oerder et al., 2015]] ; [[#Sylla--2019|Sylla et al., 2019]] ). Direct estimation of oceanic upward transport ( [[#Oyarzún--2019|Oyarzún and Brierley, 2019]] ; [[#Sylla--2019|Sylla et al., 2019]] ) and nutrient flux into the euphotic layer ( [[#Jacox--2018|Jacox et al., 2018]] ) provide a meaningful estimator of upwelling, integrating all relevant processes, including changes in wind stress curl. However, there is ''limited evidence'' from vertical velocity of climate models and missing processes in coarse-resolution climate models that presently limit this approach. Change in upper-ocean stratification ( [[#9.2.1.3|Section 9.2.1.3]] ) is projected to increase confinement of upwelling vertical velocities to near the ocean surface ( ''high confidence'' ) ( [[#Oerder--2015|Oerder et al., 2015]] ; [[#Oyarzún--2019|Oyarzún and Brierley, 2019]] ).

In summary, SROCC and this Report conclude that the California Current system has experienced some upwelling-favourable wind intensification since the 1980s ( ''high confidence'' ), while ''low agreement'' among reported wind changes in the Benguela, Canary, and Humboldt systems prevent a similar assessment. As in SROCC, there is ''low confidence'' in attribution of observed changes to anthropogenic or natural causes. New evidence reinforces our confidence in SROCC assessment that, under increased radiative forcing, EBUS winds will change with a dipole spatial pattern within each EBUS of reduction (weaker and/or shorter) at low latitude, and enhancement (stronger and/or longer) at high latitude ( ''high confidence'' ). There is ''medium confidence'' that, across all scenarios, upwelling wind changes in EBUS will remain moderate in the 21st century, within ±10–20% from present-day values.

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<span id="coastal-systems-and-marginal-seas"></span>
==== 9.2.3.6 Coastal Systems and Marginal Seas ====

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Beyond the world’s coastlines lie the shoreline, shallow estuaries, continental shelves, and deeper fjords and slopes, where depths increase rapidly from the shelves to the deep-ocean floor. It is more difficult to transport fluid across (rather than along) the shelf-break or slope ( [[#Brink--2016|Brink, 2016]] ), and estuaries and shelves have complex circulations and mixing, leading to indirect connections between the inner shelves and coastlines and offshore conditions. Coastal processes link to large-scale metrics of climate and regional effects, from changing rivers and estuaries, melt and runoff to deep water, to how changes offshore affect regional and coastal conditions.

Shelf-deep ocean exchanges involve eddying, tidal, or turbulent motions and small-scale topography such as submarine canyons; high-resolution observations and models are needed to capture these effects ( [[#Greenberg--2007|Greenberg et al., 2007]] ; [[#Capet--2008|Capet et al., 2008]] ; [[#Allen--2009|Allen and Durrieu de Madron, 2009]] ; [[#Colas--2012|Colas et al., 2012]] ; [[#Trotta--2017|Trotta et al., 2017]] ). Example coastal processes that introduce uncertainty into large-scale projections are exchange of CDW across the Antarctic shelf-break, which affects AABW formation and Antarctic ice-shelf–ocean interaction (Sections 9.2.2.3 and 9.2.3.2; [[#Stewart--2013|Stewart and Thompson, 2013]] , 2015), river and estuarine plumes and their responses to water level and hydrology change ( [[#Banas--2009|Banas et al., 2009]] ; [[#Sun--2017|Sun et al., 2017]] ), fjord dynamics linked to glacial outflows ( [[#Straneo--2015|Straneo and Cenedese, 2015]] ; [[#Torsvik--2019|Torsvik et al., 2019]] ), and changing formation of water masses in marginal seas ( [[#Kim--2001|Kim et al., 2001]] ; [[#Greene--2007|Greene and Pershing, 2007]] ; [[#Giorgi--2008|Giorgi and Lionello, 2008]] ; [[#Renner--2009|Renner et al., 2009]] ). Downscaling projections to the local level allows process detail ( [[#Foreman--2014|Foreman et al., 2014]] ; [[#Mathis--2014|Mathis and Pohlmann, 2014]] ; [[#Meier--2015|Meier, 2015]] ; [[#Tinker--2016|Tinker et al., 2016]] ). Some processes can only be simulated when coastal models are forced by larger-scale models of the atmosphere, cryosphere, or hydrosphere ( [[#Seo--2007|Seo et al., 2007]] , 2008; [[#Somot--2008|Somot et al., 2008]] ; [[#Oerder--2015|Oerder et al., 2015]] ; [[#Renault--2016|Renault et al., 2016]] ; Y. [[#Zhang--2016|]] [[#Zhang--2016|]] [[#Zhang--2016|Zhang et al., 2016]] ; [[#Wåhlin--2020|Wåhlin et al., 2020]] ), including the addition of tides ( [[#Janeković--2012|Janeković and Powell, 2012]] ; [[#Timko--2013|Timko et al., 2013]] ; [[#Tinker--2015|Tinker et al., 2015]] ; [[#Pickering--2017|Pickering et al., 2017]] ; [[#Hausmann--2020|Hausmann et al., 2020]] ). Due to coastal process complexity and small scale, linking the effects of coastal ocean changes to global ocean changes requires high-resolution modelling ( [[#Holt--2017|Holt et al., 2017]] , 2018), two-way nesting, or local mesh refinement ( [[#Fringer--2006|Fringer et al., 2006]] ; [[#Zhang--2008|Zhang and Baptista, 2008]] ; [[#Mason--2010|Mason et al., 2010]] ; [[#Dietrich--2012|Dietrich et al., 2012]] ; [[#Hellmer--2012|Hellmer et al., 2012]] ; [[#Ringler--2013|Ringler et al., 2013]] ; Q. [[#Wang--2014|]] [[#Wang--2014|Wang et al., 2014]] ; [[#Zängl--2015|Zängl et al., 2015]] ; Y.J. [[#Zhang--2016|]] [[#Zhang--2016|]] [[#Zhang--2016|Zhang et al., 2016]] ; [[#Soto-Navarro--2020|Soto-Navarro et al., 2020]] ). Coarse climate models and HighResMIP models do not represent some coastal phenomena such as cross-shelf exchanges and sub-mesoscale eddies, which require 1 km or finer resolution. Thus, there is ''low confidence'' in projecting centennial scale coastal climate change where regional downscaling or refinement is lacking. There is ''high confidence'' in the ability of regional coupled models to improve coastal climate change process understanding and provide regional information ( [[IPCC:Wg1:Chapter:Chapter-12#12.4|Section 12.4]] ), but many sites globally await such projections.

<div id="9.2.4" class="h2-container"></div>

<span id="steric-and-dynamic-sea-level-change"></span>
=== 9.2.4 Steric and Dynamic Sea Level Change ===

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<span id="global-mean-thermosteric-sea-level-change"></span>
==== 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'' ).

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

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