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

=== 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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'''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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'''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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'''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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'''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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==== 9.2.2.2 Ocean Salinity ====

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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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==== 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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