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

=== 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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'''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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==== 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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==== 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.

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